Innovative andApplied Research on
Platinum-Group andRare Earth Elements
Dedicated to the Work andMemory of Dr. Demetrios G.
Eliopoulos, IGME (Greece)

Edited by

Maria Economou-Eliopoulos, Federica Zaccarini and
Giorgio Garuti
Printed Edition of the Special Issue Published in Minerals

Innovative and Applied Researchon Platinum-Group and Rare
Earth Elements

Innovative and Applied Researchon Platinum-Group and Rare
Earth Elements
Dedicated to the Work and Memory of
Dr. Demetrios G. Eliopoulos, IGME (Greece)

Special Issue Editors
Maria Economou-Eliopoulos
Federica Zaccarini
Giorgio Garuti

MDPI • Basel • Beijing • Wuhan • Barcelona • Belgrade • Manchester • Tokyo • Cluj • Tianjin

Special Issue Editors
Maria Economou-Eliopoulos

Federica Zaccarini

Giorgio Garuti

University of Athens

University of Leoben

University of Leoben

GreeceAustria

Austria

Editorial Ofﬁce
MDPI
St. Alban-Anlage 664052 Basel, Switzerland

This is a reprint of articles from the Special Issue published online in the open access journal
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PGE REE).

For citation purposes, cite each article independently as indicated on the article page online and as
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Cover image courtesy of Federica Zaccarini.

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Contents
About the Special Issue Editors . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . vii
Maria Economou-Eliopoulos, Federica Zaccarini and Giorgio Garuti
Editorial for the Special Issue “Innovative and Applied Research on Platinum-Group and Rare
Earth Elements”
Reprinted from: Minerals 2020, 10, 493, doi:10.3390/min10060493 . . . . . . . . . . . . . . . . . .

1

Luca Bindi, Federica Zaccarini, Paola Bonazzi, Tassos Grammatikopoulos,
Basilios Tsikouras, Chris Stanley and Giorgio Garuti
Eliopoulosite, V7S8, A New Sulfide from the Podiform Chromitite of the Othrys
Ophiolite, Greece
Reprinted from: Minerals 2020, 10, 245, doi:10.3390/min10030245 . . . . . . . . . . . . . . . . . .

5Luca Bindi, Federica Zaccarini, Elena Ifandi, Basilios Tsikouras, Chris Stanley,
Giorgio Garuti and Daniela Mauro
Grammatikopoulosite, NiVP, a New Phosphide from the Chromitite of the Othrys
Ophiolite, Greece
Reprinted from: Minerals 2020, 10, 131, doi:10.3390/min10020131 . . . . . . . . . . . . . . . . . . 19
Olga N. Kiseleva, Evgeniya V. Airiyants, Dmitriy K. Belyanin and Sergey M. Zhmodik
Podiform Chromitites and PGE Mineralization in the Ulan-Sar’dag Ophiolite(East Sayan, Russia)
Reprinted from: Minerals 2020, 10, 141, doi:10.3390/min10020141 . . . . . . . . . . . . . . . . . . 31
Raisa G. Kravtsova, Vladimir L. Tauson, Artem S. Makshakov, Nikolay V. Bryansky and
Nikolay V. Smagunov
Platinum Group Elements in Arsenopyrites and Pyrites of the Natalkinskoe Gold Deposit(Northeastern Russia)
Reprinted from: Minerals 2020, 10, 318, doi:10.3390/min10040318 . . . . . . . . . . . . . . . . . . 61
Federica Zaccarini and Giorgio Garuti
Zoned Laurite from the Merensky Reef, Bushveld Complex, South Africa: “Hydrothermal”
in Origin?
Reprinted from: Minerals 2020, 10, 373, doi:10.3390/min10040373 . . . . . . . . . . . . . . . . . . 97
Martiya Sadeghi, Nikolaos Arvanitidis and Anna Ladenberger
Geochemistry of Rare Earth Elements in Bedrock and Till, Applied in the Context of Mineral
Potential in Sweden
Reprinted from: Minerals 2020, 10, 365, doi:10.3390/min10040365 . . . . . . . . . . . . . . . . . . 111
Christina Stouraiti, Vassiliki Angelatou, Sofia Petushok, Konstantinos Soukis and
Demetrios Eliopoulos
Effect of Mineralogy on the Beneficiation of REE from Heavy Mineral Sands: The Case of Nea
Peramos, Kavala, Northern Greece
Reprinted from: Minerals 2020, 10, 387, doi:10.3390/min10050387 . . . . . . . . . . . . . . . . . . 135
Maria Economou-Eliopoulos, Robert Frei and Ioannis Mitsis
Factors Controlling the Chromium Isotope Compositions in Podiform Chromitites
Reprinted from: Minerals 2020, 10, 10, doi:10.3390/min10010010 . . . . . . . . . . . . . . . . . . . 153v

Ioannis-Porfyrios D. Eliopoulos and George D. Eliopoulos
Factors Controlling the Gallium Preference in High-Al Chromitites
Reprinted from: Minerals 2019, 9, 623, doi:10.3390/min9100623 . . . . . . . . . . . . . . . . . . . . 175
Demetrios G. Eliopoulos and Maria Economou-Eliopoulos
Trace Element Distribution in Magnetite Separates of Varying Origin: Genetic andExploration SignificanceReprinted from: Minerals 2019, 9, 759, doi:10.3390/min9120759 . . . . . . . . . . . . . . . . . . . . 191
Demetrios G. Eliopoulos, Maria Economou-Eliopoulos, George Economou andVassilis Skounakis
Mineralogical and Geochemical Constraints on the Origin of Mafic–Ultramafic-Hosted
Sulphides: The Pindos Ophiolite Complex
Reprinted from: Minerals 2020, 10, 454, doi:10.3390/min10050454 . . . . . . . . . . . . . . . . . . 219
Connor A. Gray and Adrian D. Van Rythoven
A Comparative Study of Porphyry-Type Copper Deposit Mineralogies by Portable X-ray
Fluorescence and Optical Petrography
Reprinted from: Minerals 2020, 10, 431, doi:10.3390/min10050431 . . . . . . . . . . . . . . . . . . 239

viAbout the Special Issue Editors
Maria Economou-Eliopoulos Geologist and Chemist, was a Professor (1993–2014) of Economic
Geology and Geochemistry and Director of the Postgraduate programme “Applied Environmental
Geology” (2000–2014), and is currently Professor Emeritus at the National and Kapodistrian
University of Athens.

Her scientiﬁc interests include geochemistry of ore deposits, from theore-forming processes to the environmental problems. Her post-doctoral research at the Universityof Toronto, Imperial College, U.S. Geological Survey (Reston) and Universities of Europe wasmainly focused on the platinum-group elements in maﬁc-ultramaﬁc rocks and associated deposits(chromite, sulﬁdes, magnetite, and laterites), and porphyry Cu-Au±Pd±Pt systems, which are her
main research topic currently. Maria is the author of more than 150 publications in scientiﬁc journals,
thematic issues, chapters in books, monographs, a fellow of the SEG (Society-Economic-Geologists),
evaluator of research proposals (national and international), a reviewer of scientiﬁc papers in morethan 30 leading geological journals, has over 2500 citations, and an h-index of 29. Her distinctions
include awards from the National Academy of Athens (2008, 2001, 1991, 1986, and 1981) for the
discovery of a new mineral (Theophrastite) and palladium in a porphyry-Cu deposit of Greece andconstraints on the genesis of ore deposits.
Federica Zaccarini from Modena, Italy obtained an Italian-European habilitation as a fullProfessor in Mineralogy, Petrology, Geochemistry, Volcanology, and Ore Deposits. She is a Senior
Scientist at the University of Leoben (Austria) and Head of the E.F. Stumpﬂ Electron Laboratory.
Federica has received a large number of international awards including the following:• 1998 Ph.D. Student grant, Society of Economic Geology (SEG), USA• 2001 Johndino Nogara, Italian Society of Mineralogy and Petrology, ITALY• 2004–2006 FWF Lise Meitner fellowship, University of Leoben, AUSTRIA• 2005 Johndino Nogara, Italian Society of Mineralogy and Petrology, ITALY• 2006 Juan de la Cierva Fellowship, University of Granada, SPAIN• 2007 Ramon y Cajal fellowship, SPAIN.
Her main research interests include the characterization of opaque mineral species, particularly
using optical and electron microscopes, and electron microprobe analysis, investigation andcharacterization of accessory minerals, analytical techniques for the noble metals and theirgeochemistry, mineralogy, and origin of ore deposits related with maﬁc-ultramaﬁc rocks, platinumgroup minerals in chromitites from ophiolite, stratiform complexes and Ural-Alaskan type intrusion,and sulﬁdes deposits and sulﬁde deposits in ophiolitic complexes. Her work has been recognized by
the international scientiﬁc community and the International Mineralogical Association has adopted
a new mineral name in her honor (Zaccarinite). Federica is the author of more than 100 papers andseveral book chapters. She has over 2000 citations and an h-index of 24.

vii

Giorgio Garuti from Modena, Italy was a Professor of Ore Deposits at the University of Modena
from 1982 to 2009. Currently, he is a lecturer at the University of Leoben (Austria). Giorgio is aninternationally recognized expert in petrology, geochemistry, and mineralogy of maﬁc-ultramaﬁcrocks and associated ore deposits. His main research interests include the mineralogy andgeochemistry of noble metals, especially platinum group elements and their deposits, such as
magmatic sulﬁdes and chromitites. Due to his experience, Giorgio worked intensively in several areas
worldwide (North–Central–South America, Africa, Russia, Europe, and Asia). His current research
includes the investigation of ore deposits related to ophiolites, with a special focus on the presence ofgold and uranium minerals. He was the Associate Editor of the Canadian Mineralogist and presently
he is a member of the Editorial Board of the journals Geologica Acta and Oﬁoliti. His work has beenrecognized by the international scientiﬁc community and the International Mineralogical Association
has adopted a new mineral name in his honor (Garutite). Giorgio is the author of more than 100
papers and several book chapters. He has over 2000 citations and an h-index of 29.

viiiminerals
Editorial

Editorial for the Special Issue “Innovative andApplied Research on Platinum-Group and Rare
Earth Elements”
Maria Economou-Eliopoulos 1, *, Federica Zaccarini 2 and Giorgio Garuti 2
1
2*

Department of Geology and Geoenvironment, University of Athens, 15784 Athens, Greece
Department of Applied Geological Sciences and Geophysics, University of Leoben, A-8700 Leoben, Austria;
federica.zaccarini@unileoben.ac.at (F.Z.); giorgio.garuti1945@gmail.com (G.G.)
Correspondence: econom@geol.uoa.gr

Received: 19 May 2020; Accepted: 27 May 2020; Published: 28 May 2020

This Special Issue “Innovative and Applied Research on Platinum-group and Rare Earth Elements”
is dedicated to the work and memory of Demetrios Eliopoulos, IGME (Institute of Geology and Mineral
Exploration), Greece who passed away on 19 April 2019.
Demetrios Eliopoulos (1947–2019) started his research work (nearly 40 years ago) on base andprecious metals with his Master’s thesis (Western University, Canada) and Ph.D thesis (Southampton,
UK). Over the course of his career at IGME (Greek Institute of Geology and Mineral Exploration),
Demetrios was the IGME representative of R&D, European Commission Projects, Head of Economic
Geology, and a contributor to numerous fundamental research reports and publications. His main
distinctions include: his membership on the Editorial Board of Mineralium Deposita, being a Fellow of
the SEG (Society-Economic-Geologists) and SGA (Society for Geology Applied to Mineral Deposits),
and the Award from the Academy of Athens (discovery of palladium in a porphyry-Cu deposit ofGreece). Demetrios was the chairman of the Organizing Committee of the 7th Biennial SGA Meeting,held in Athens, Greece, August 24–28, 2003. He was involved in fruitful scientiﬁc collaboration over
the years with his wife and colleague Maria Economou-Eliopoulos. Their work was recognized by theinternational community, and the IMA has adopted a new mineral name in honor of Demetrios andMaria (Eliopoulosite, IMA 2019-96, this volume). This Special Issue contains a series of papers, including
two reporting on the ﬁnding of new minerals—Grammatikopoulosite, NiVP, a new phosphide, andEliopoulosite, V7 S8 , a new pyrrhotite like mineral—both from chromitites of the Othrys ophiolitecomplex, approved by the Commission of New Minerals, Nomenclature, and Classiﬁcation of theInternational Mineralogical Association (IMA 2019-090 and 2019-96) [1,2]. The geological environment,mode of the occurrence, physical and chemical properties, and crystallographic data are provided.
Problems concerning the distribution, mineralogy and ﬁeld relationships of PGE-enriched ores, which
are important for our understanding of the metallogenic controls on the concentration of PGE in
ophiolite complexes and their exploration, are addressed in three papers. Kiseleva et al. [3] present theﬁrst detailed data on the chromitites and PGE mineralization in the Ulan-Sar’dag ophiolite locatedin the Central Asian Fold Belt (East Sayan, Russia), and propose a scheme for the sequence of theformation and transformation of the PGM at various stages of the evolution of that ophiolite complex.
Kravtsova et al. [4] present the peculiarities of the distribution and binding forms of PGE in thearsenopyrites and pyrites of the Natalkinskoe gold ore deposit in NE Russia. They conclude thatthe possibility of their extraction from that type of ore may signiﬁcantly increase the quality andvalue of the extracted materials at the Natalkinskoe deposit. Zaccarini and Garuti [5] describe theoccurrence of zoned laurite found in the Merensky Reef of the Bushveld layered intrusion, South
Africa. Speciﬁcally, zoned laurites from the Merensky Reef are characterized by textural position,
composition, and zoning—suggesting that they are “hydrothermal” in origin, having crystallized in the

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Minerals 2020, 10, 493

presence of a Cl- and As-rich hydrous solution, at temperatures much lower than those typical of the
precipitation of magmatic laurite. The paper by Sadeghi et al. [6] deals with the study of the rare earth
elements (REE) in till and bedrock, as well as in mineral deposits associated with apatite-iron oxidemineralizations, various skarn deposits, hydrothermal deposits, and alkaline-carbonatite intrusionsin Sweden. In this study, analytical data of samples collected from REE mineralizations during the
EURARE (www.eurare.org) project are compared with bedrock and till REE geochemistry—both
sourced from databases available at the Geological Survey of Sweden (www.sgu.se). These results
are useful in the assessment of REE mineral potential in areas where REE mineralizations are poorly
explored or even undiscovered. The paper by Stouraiti et al. [7] presents the eﬀect of mineralogy on
the beneﬁciation of REE from heavy mineral sands from Nea Peramos, Kavala, northern Greece. They
provide the characterization of the fractions separated by a magnetic separator using X-ray powder
diﬀraction (XRD) and bulk Inductively Coupled Plasma and Mass Spectrometry (ICP-MS) chemical
analyses, and show that at high intensities of magnetic separation, they were strongly enriched inLight Rare Earth Elements (LREE) compared to the non-magnetic fraction. This new information cancontribute to the optimization of the beneﬁciation process, that can be applied for REE recovery from
Heavy Mineral (HM) black sands. A group of papers deal with the genetic signiﬁcance of a spectrumof trace elements and stable chromium isotopes in chromitites. The paper by Economou-Eliopoulos
et al. [8] on the factors controlling the chromium isotope compositions in podiform chromitites is anapplication of Cr stable isotope (δ53 Cr values) compositions to the investigation of magmatic andpost-magmatic eﬀects on chromitites associated with ophiolite complexes. Chromitites from the Balkan
Peninsula depict a wide range of δ53 Cr values. Signatures range from positively to slightly negatively
fractionated δ53 Cr values—even in individual relatively large deposits. Positively fractionated δ53 Cr
values of all chromitite samples from Othrys and of high-Al chromitites from Skyros, coupled with anegatively correlated trend between δ53 Cr and Cr/(Cr + Al), may reﬂect the control of δ53 Cr by degreeof partial melting and by magma fractionation. This is best exempliﬁed by high-Al chromitites from thecumulate sequence of the Vourinos complex. The paper by Eliopoulos and Eliopoulos [9] on the factors
controlling the gallium preference in high-Al chromitites presents geochemical—including Ga—andmineral chemistry data on chromitites associated with the ophiolite complexes of Greece and elsewhere.
Potential factors controlling the lower Ga contents in high-Cr chromitites compared to high-Al ones are
the composition of the parent magma, temperature, redox conditions, the disorder degree of spinels,
and the ability of Al3+ to occupy both octahedral and tetrahedral sites, in contrast to the competing
Cr3+ that can occupy only octahedral sites (due to its electronic conﬁguration) while the Ga3+ shows astrong preference on tetrahedral sites. The paper by Eliopoulos and Economou-Eliopoulos [10] on the
trace element distribution in magnetite separates of varying origin and genetic/exploration signiﬁcanceprovides SEM/EDS and ICP-MS data on representative magnetite samples covering various geotectonic
settings and rock-types, such as calc-alkaline and ophiolitic rocks, porphyry-Cu deposits, skarn-type,ultramaﬁc lavas, black coastal sands, and metamorphosed Fe–Ni-laterites deposits. They concluded
that, despite the potential re-distribution of trace elements—including Rare Earth Elements (REE)and Platinum Group Elements (PGE) in magnetite-bearing deposits—the data may provide valuableevidence for their origin and exploration. In particular, the trace element content and the presence ofabundant magnetite provide valuable evidence for discrimination between Cu–Au–Pd–Pt- porphyrysystems and those lacking precious metals. The paper by Gray and Van Rythoven [11] deals with acomparative study between porphyry-Cu deposits of the Western Cordillera, using portable X-ray
ﬂuorescence analysis (pXRF) and optical microscopy. Although the proposed research methodology
cannot reach the levels of precision and accuracy required to meet standards for mineral resources, thecombined use of pXRF in the ﬁeld with optical microscopy provides a fast and cost-eﬀective method forthe exploration of large unexplored porphyry intrusions characterized by extensive haloes of alteration.
The application of the proposed methodology to porphyry intrusions can inﬂuence decisions regarding
geological mapping, cross-sections, and the selection of samples for a more detailed petrographic andgeochemical investigation of porphyry-type deposits, as well as inﬂuence exploration for potential

2

Minerals 2020, 10, 493

sources of Cu, Mo and precious metals (Au and platinum-group elements). The paper by Eliopoulos
et al. [12] is focused on sulphides of both Cyprus-type and Fe-Cu-Co-Zn sulphides associated withmagnetite within gabbro, close to its tectonic contact with serpentinized harzburgite of the Pindos
ophiolite complex. The higher Zn, Se, Mo, Au, Ag, Hg, and Sb contents and lower Ni contents in the
Pindos compared to the Othrys sulphides may reﬂect inheritance of a primary magmatic signature.
Relatively high Cu/(Cu+Ni) and Pt/(Pt+Pd), and low Ni/Co ratios for sulphides from the deeper part ofthe complex suggest either no magmatic origin, or a complete transformation of preexisting magmatic
assemblages. However, the recorded textural and mineralogical features of both sulphide types from
Pindos resemble the Fe-Cu-Zn-Co-Ni mineralization reported in other maﬁc–ultramaﬁc-hosted VMS
deposits, which have shown mineralogical, compositional and textural analogies with present-day
counterparts on ultramaﬁc-rich substrate seaﬂoor VMS deposits formed by hydrothermal processes.
Author Contributions: M.E.-E., F.Z. and G.G. wrote the editorial. All authors have read and agreed to the
published version of the manuscript.
Conﬂicts of Interest: The authors declare no conﬂict of interest.

References
1.

Bindi, L.; Zaccarini, F.; Ifandi, E.; Tsikouras, B.; Stanley, G.; Garuti, G.; Mauro, D. Grammatikopoulosite,
NiVP, a New Phosphide from the Chromitite of the Othrys Ophiolite, Greece. Minerals 2020, 10, 131.
2.
Bindi, L.; Zaccarini, F.; Bonazzi, P.; Grammatikopoulos, T.; Tsikouras, B.; Stanley, C.; Garuti, G. Eliopoulosite,
V7 S8 , A New Sulﬁde from the Podiform Chromitite of the Othrys Ophiolite, Greece. Minerals 2020, 10, 245.
[CrossRef]
3.
Kiseleva, O.N.; Airiyants, E.V.; Belyanin, D.K.; Zhmodik, S.M. Podiform Chromitites and PGE Mineralizationin the Ulan-Sar’dag Ophiolite (East Sayan, Russia). Minerals 2020, 10, 141. [CrossRef]
4.
Kravtsova, R.G.; Tauson, V.L.; Makshakov, A.S.; Bryansky, N.V.; Smagunov, N.V. Platinum Group Elements
in Arsenopyrites and Pyrites of the Natalkinskoe Gold Deposit (Northeastern Russia). Minerals 2020, 10, 318.
5.
Zaccarini, F.; Garuti, G. Zoned Laurite from the Merensky Reef, Bushveld Complex, South Africa:“Hydrothermal” in Origin? Minerals 2020, 10, 373. [CrossRef]
6.
Sadeghi, M.; Arvanitidis, N.; Ladenberger, A. Geochemistry of Rare Earth Elements in Bedrock and Till,
Applied in the Context of Mineral Potential in Sweden. Minerals 2020, 10, 365. [CrossRef]
7.
Stouraiti, C.; Angelatou, V.; Petushok, S.; Soukis, K.; Eliopoulos, D. Eﬀect of mineralogy on the beneﬁciationof REEs from heavy mineral sands: The case of Nea Peramos, Kavala, northern Greece. Minerals 2020, 10,387. [CrossRef]
Economou-Eliopoulos, M.; Frei, R.; Ioannis, I. Factors Controlling the Chromium Isotope Compositions in
8.
Podiform Chromitites. Minerals 2020, 10, 10. [CrossRef]
9.
Eliopoulos, I.P.D.; Eliopoulos, G.D. Factors Controlling the Gallium Preference in High-Al Chromitites.
Minerals 2020, 9, 623. [CrossRef]
10. Eliopoulos, D.G.; Economou-Eliopoulos, M. Trace Element Distribution in Magnetite Separates of Varying
Origin: Genetic and Exploration Signiﬁcance. Minerals 2020, 9, 759. [CrossRef]
11. Gray, C.A.; Van Rythoven, A.D. A Comparative Study of Porphyry-Type Copper Deposit Mineralogies by
Portable X-ray Fluorescence and Optical Petrography. Minerals 2020, 10, 431. [CrossRef]
12. Eliopoulos, D.G.; Economou-Eliopoulos, M.; Economou, G.; Skounakis, V. Mineralogical and Geochemical
Constraints on the Origin of Maﬁc–Ultramaﬁc-Hosted Sulphides: The Pindos Ophiolite Complex. Minerals
2020, 10, 454.
© 2020 by the authors. Licensee MDPI, Basel, Switzerland. This article is an open accessarticle distributed under the terms and conditions of the Creative Commons Attribution(CC BY) license (http://creativecommons.org/licenses/by/4.0/).

3

minerals
ArticleEliopoulosite, V7S8, A New Sulﬁde from the
Podiform Chromitite of the Othrys Ophiolite, Greece
Luca Bindi 1, *, Federica Zaccarini 2 , Paola Bonazzi 1 , Tassos Grammatikopoulos 3 ,Basilios Tsikouras 4 , Chris Stanley 5 and Giorgio Garuti 2
1
2
3
4
5

*

Dipartimento di Scienze della Terra, Università degli Studi di Firenze, I-50121 Florence, Italy;
paola.bonazzi@uniﬁ.it
Department of Applied Geological Sciences and Geophysics, University of Leoben, A-8700 Leoben, Austria;
federica.zaccarini@unileoben.ac.at (F.Z.); giorgio.garuti1945@gmail.com (G.G.)SGS Canada Inc., 185 Concession Street, PO 4300, Lakeﬁeld, ON K0L 2H0, Canada;
Tassos.Grammatikopoulos@sgs.com
Faculty of Science, Physical and Geological Sciences, Universiti Brunei Darussalam, BE 1410 Gadong,
Brunei Darussalam; basilios.tsikouras@ubd.edu.bn
Department of Earth Sciences, Natural History Museum, London SW7 5BD, UK; c.stanley@nhm.ac.uk
Correspondence: luca.bindi@uniﬁ.it; Tel.: +39-055-275-7532

Received: 22 February 2020; Accepted: 6 March 2020; Published: 8 March 2020

Abstract: The new mineral species, eliopoulosite, V7 S8 , was discovered in the abandoned chromiummine of Agios Stefanos of the Othrys ophiolite, located in central Greece. The investigated samples
consist of massive chromitite hosted in a strongly altered mantle tectonite, and are associated withnickelphosphide, awaruite, tsikourasite, and grammatikopoulosite. Eliopoulosite is brittle andhas a metallic luster. In plane-reﬂected polarized light, it is grayish-brown and shows no internal
reﬂections, bireﬂectance, and pleochroism. It is weakly anisotropic, with colors varying from lightto dark greenish. Reﬂectance values of mineral in air (Ro , Re’ in %) are: 34.8–35.7 at 470 nm, 38–39
at 546 nm, 40–41.3 at 589 nm, and 42.5–44.2 at 650 nm. Electron-microprobe analyses yielded amean composition (wt.%) of: S 41.78, V 54.11, Ni 1.71, Fe 1.1, Co 0.67, and Mo 0.66, totali 100.03.
On the basis of Σatoms = 15 apfu and taking into account the structural data, the empirical formulaof eliopoulosite is (V6.55 Ni0.19 Fe0.12 Co0.07 Mo0.04 )Σ = 6.97 S8.03 . The simpliﬁed formula is (V, Ni, Fe)7 S8and the ideal formula is V7 S8 , which corresponds to V 58.16%, S 41.84%, total 100 wt.%. The density,
based on the empirical formula and unit-cell volume reﬁned form single-crystal structure XRD data,is 4.545 g·cm−3 . The mineral is trigonal, space group P32 21, with a = 6.689(3) Å, c = 17.403(6) Å,V = 674.4(5) Å3 , Z = 3, and exhibits a twelve-fold superstructure (2a × 2a × 3c) of the NiAs-type
subcell with V-atoms octahedrally coordinated by S atoms. The distribution of vacancies is discussed
in relation to other pyrrhotite-like compounds. The mineral name is for Dr. Demetrios Eliopoulos(1947–2019), a geoscientist at the Institute of Geology and Mineral Exploration (IGME) of Greece andhis widow, Prof. Maria Eliopoulos (nee Economou, 1947), University of Athens, Greece, for theircontributions to the knowledge of ore deposits of Greece and to the mineralogical, petrographic,
and geochemical studies of ophiolites, including the Othrys complex. The mineral and its namehave been approved by the Commission of New Minerals, Nomenclature, and Classiﬁcation of theInternational Mineralogical Association (No. 2019-96).
Keywords: eliopoulosite; sulﬁde; chromitite; Agios Stefanos mine; Othrys; ophiolite; Greece

1. Introduction
Only eight minerals containing V and S are in the list of valid species approved by the International
Mineralogical Association (IMA). They include, in alphabetic order: colimaite, K3 VS4 [1]; colusite,
Minerals 2020, 10, 245; doi:10.3390/min10030245

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Minerals 2020, 10, 245

Cu13 VAs3 S16 [2]; germanocolusite, Cu13 VGe3 S16 [3]; merelaniite, Mo4 Pb4 VSbS15 [4]; nekrasovite,
Cu13 VSn3 S16 [5]; patronite, VS4 [6]; stibiocolusite, Cu13 V(Sb,Sn,As)3 S16 [7]; and sulvanite, Cu3 VS4 [8].
All of them, with the exception of patronite, a mineral discovered in 1906 in the Minas Ragra ofPeru [6], are sulﬁdes or sulfosalts characterized by a complex composition including other metals
besides vanadium.
During a recent investigation of the heavy mineral concentrates from a chromitite collected in the
Othrys ophiolite (central Greece), three new minerals were discovered. Two of them are phosphides,
namely tsikourasite, Mo3 Ni2 P1 + x [9] and grammatikopoulosite, NiVP [10]. A chemical and structural
study revealed the third mineral to be a new sulﬁde, trigonal in symmetry and having the ideal formulaV7 S8 . The mineral and its name were approved by the Commission of New Minerals, Nomenclature
and Classiﬁcation of the International Mineralogical Association (No. 2019-096). The new mineral
has been named after Dr. Demetrios Eliopoulos (1947–2019), a geoscientist at the Institute of Geology
and Mineral Exploration (IGME) of Greece and his widow, Prof. Maria Eliopoulos (nee Economou,1947), University of Athens, Greece, for their contributions to the knowledge of ore deposits of Greece
and to the mineralogical, petrographic, and geochemical studies of ophiolites, including the Othrys
complex. Holotype material is deposited in the Mineralogical Collection of the Museo di Storia
Naturale, Università di Pisa, Via Roma 79, Calci (Pisa, Italy), under catalogue number 19911 (same typespecimen of grammatikopoulosite).
2. Geological Background and Occurrence of Eliopoulosite
Eliopoulosite was discovered in a heavy-mineral concentrate obtained from massive chromitite
collected in the mantle sequence of the Mesozoic Othrys ophiolite, located in central Greece (Figure 1A).
The geology, petrography, and geodynamic setting of the Otrhys ophiolite have been discussed byseveral authors [11–24]. The studied sample was collected in the abandoned chromium mine of Agios
Stefanos, located a few km southwest of the Domokos village (Figure 1B). In the studied area, a mantle
tectonite composed of plagioclase lherzolite, harzburgite, and minor harzburgite-dunite occurs incontact with rocks of the crustal sequence (Figure 1C). The investigated chromitite was collected fromthe harzburgite-dunite that is cut across by several dykes of gabbro (Figure 1C).
The ophiolite of Othrys is a complete but dismembered suite (Mirna Group) and consists of threestructural units: the uppermost succession with variably serpentinized peridotites, which is structurallybounded by an ophiolite mélange; the intermediate Kournovon dolerite, including cumulate gabbro
and local rhyolite; and the lower Sipetorrema Pillow Lava unit also including basaltic ﬂows, siltstones,
and chert. The Mirna Group constitutes multiple inverted thrust sheets, which were eventually
obducted onto the Pelagonian Zone, during the Late Jurassic-Early Cretaceous [11–14]. Three typesof basalts with diﬀerent geochemical signatures have been described: i) alkaline within-plate (WPB),
occurring in the mélange, which is related to oceanic seamounts or to ocean-continent transition zones;ii) normal-type mid-ocean ridge (N-MORB); and iii) low-K tholeiitic (L-KT) rocks, which are erupted
close to the rifted margin of an ocean-continent transition zone. Radiolarian data from interbedded
cherts suggest Middle to Late Triassic ages [15]. The Othrys ophiolite is structurally divided into thewest and east Othrys suites, which are thought to derive from diﬀerent geotectonic environments.
The former is related to an extension regime (back-arc basin or MORB [16–18]), while the latter
formed in a supra-subduction zone (SSZ) setting [17–19]. This diﬀerence is reﬂected in the contrastinggeochemical compositions of their ultramaﬁc rocks, as well as in the diverse platinum group minerals(PGM) assemblages of their host chromitites [20,21], which suggest large mantle heterogeneities in thisarea. Conﬂicting views with regards to the evolution of the Othrys ophiolite suggest the involvementof Pindos [11,22,23] or Vardar (Axios) [15] oceanic domains.
The studied spinel-supergroup minerals that host eliopoulosite can been classiﬁed asmagnesiochromite, although their composition is rather heterogeneous [9,10,25] with the amounts of
Cr2 O3 (44.96–51.64 wt.%), Al2 O3 (14.18–20.78 wt.%), MgO (13.34–16.84 wt.%), and FeO (8.3–13.31 wt.%)
varying signiﬁcantly through the sampled rock. The calculated Fe2 O3 is relatively high, and ranges

6Minerals 2020, 10, 245

from 6.72 to 9.26 wt.%. The amounts of MnO (0.33–0.60 wt.%), V2 O3 (0.04–0.30 wt.%), ZnO (up to 0.07wt.%), and NiO (0.03–0.24 wt.%) are low. The maximum TiO2 content is 0.23 wt.%, which is typical forthe podiform chromitites hosted in the mantle sequence of ophiolite complexes.

Figure 1. Location of the Othrys complex in Greece (A); general geological map of the Othrys ophiolite
showing the location of the Agios Stefanos chromium mine (B) and (C) detailed geological setting of
the Agios Stefanos area (modiﬁed after [10] and [5]).

3. Analytical Methods
The heavy minerals were concentrated at SGS Mineral Services, Canada, following themethodology described by several authors [9,10,21,25,26]. After the concentration process, the heavy
minerals were embedded in epoxy blocks and then polished for mineralogical examination.
Quantitative chemical analyses and acquisition of back-scattered electron images of eliopoulosite
were performed with a JEOL JXA-8200 electron-microprobe, installed in the E. F. Stumpﬂ laboratory,
Leoben University, Austria, operating in wavelength dispersive spectrometry (WDS) mode. Major andminor elements were determined at 20 kV accelerating voltage and 10 nA beam current, with 20 s as
the counting time for the peak and 10 s for the backgrounds. The beam diameter was about 1 μm in
size. For the WDS analyses, the following lines and diﬀracting crystals were used: S = (Kα, PETJ),
V, Ni, Fe, Co = (Kα, LIFH), and Mo = (Lα, PETJ). The following standards were selected: metallic
vanadium for V, pyrite for S and Fe, millerite for Ni, molybdenite for Mo, and skutterudite for Co.

7

Minerals 2020, 10, 245

The ZAF correction method was applied. Automatic correction was performed for interference Mo-S.
Representative analyses of eliopoulosite are listed in Table 1.
Table 1. Electron microprobe analyses (wt.%) of eliopoulosite.
Analysis

S

V

Ni

Fe

Co

Mo

Total

1
2
3
4
5
6
7
8
9
10
11
12
13
14
15
16
17
18
19
20
average

41.34
41.35
41.36
41.39
41.42
41.44
41.57
41.65
41.67
41.73
41.76
41.82
41.90
41.94
41.99
42.03
42.12
42.28
42.33
42.62
41.78

53.08
53.45
53.55
53.71
53.90
53.91
53.97
54.03
54.03
54.06
54.09
54.16
54.21
54.25
54.40
54.40
54.56
54.64
54.80
55.02
54.11

1.24
1.33
1.39
1.41
1.51
1.54
1.55
1.58
1.65
1.67
1.69
1.70
1.72
1.72
1.94
1.99
2.02
2.13
2.22
2.27
1.71

0.87
0.92
0.97
1.00
1.01
1.09
1.09
1.09
1.10
1.11
1.11
1.12
1.12
1.13
1.16
1.16
1.18
1.19
1.21
1.47
1.1

0.48
0.52
0.55
0.55
0.61
0.63
0.63
0.63
0.64
0.64
0.65
0.68
0.69
0.72
0.72
0.73
0.76
0.78
0.79
0.91
0.67

0.55
0.57
0.57
0.57
0.58
0.59
0.59
0.60
0.63
0.64
0.65
0.67
0.68
0.68
0.69
0.69
0.70
0.75
0.81
0.90
0.66

99.01
99.29
99.37
99.43
99.72
99.75
99.77
99.82
99.86
99.92
100.18100.20100.26100.27100.28100.31100.78100.78100.84100.84100.03

A small grain of eliopoulosite was hand-picked from the polished section under a reﬂected lightmicroscope. The crystal (about 80 μm in size) was carefully and repeatedly washed in acetone andmounted on a 5 μm-diameter carbon ﬁber, which was, in turn, attached to a glass rod in preparation of
the single-crystal X-ray diﬀraction measurements.
Single-crystal X-ray diﬀraction data were collected at the University of Florence (Italy) using a
Bruker D8 Venture equipped with a Photon II CCD detector, with graphite-monochromatized MoKα
radiation (λ = 0.71073 Å). Intensity data were integrated and corrected for standard Lorentz-polarization
factors with the software package Apex3 [27,28]. A total of 1289 unique reﬂections was collected up to2θ = 62.24◦ .
The reﬂectance measurements on eliopoulosite were carried out using a WTiC standard and a
J&M TIDAS diode array spectrophotometer at the Natural History Museum of London, UK.
4. Physical and Optical Properties
In the polished section, eliopoulosite occurs as tiny crystals (from 5 μm up to about 80 μm) and isanhedral to subhedral in habit. It consists of polyphase grains associated with other minerals, such as
tsikourasite, nickelphosphide, awaruite, and grammatikopoulosite (Figure 2).
In plane-reﬂected polarized light, eliopoulosite is grayish-brown and shows no internal reﬂections,bireﬂectance, and pleochroism. It is weakly anisotropic, with colors varying from light to dark greenish.
Reﬂectance values of the mineral in air (R in %) are listed in Table 2 and shown in Figure 3.

8Minerals 2020, 10, 245

Figure 2. Digital image in reflected plane polarized light (A,B) and back-scattered electron image(C,D) showing eliopoulosite from the chromitite of Agios Stefanos. Abbreviations: Elp = eliopoulosite,
Grm = grammatikopoulosite, Tsk = tsikourasite, Aw = awaruite, Npd = nickelphosphide, Chr = chromite,
Epx = epoxy.
Table 2. Reﬂectance values of eliopoulosite, those required by the Commission on Ore Mineralogy(COM) are given in bold.
λ (nm)

Ro (%)Re’ (%)

400
420
440
460
470
480
500
520
540
546
560
580
589
600
620
640
650
660
680
700

32.0
32.7
33.5
34.3
34.8
35.2
36.1
36.9
37.7
38.0
38.6
39.5
40.0
40.5
41.2
42.1
42.5
42.8
43.5
44.3

33.1
33.7
34.5
35.3
35.7
36.1
37.0
37.8
38.7
39.0
39.7
40.8
41.3
41.9
42.9
43.8
44.2
44.6
45.5
46.2

9

Minerals 2020, 10, 245

Figure 3. Reﬂectance data for eliopoulosite.

Due to the small amount of available material, density was not measured. The calculated density
was = 4.545 g·cm−3 , based on the empirical formula and the unit-cell volume reﬁned from single-crystal
XRD data.
5. X-Ray Crystallography and Chemical Composition
The mineral is trigonal, with a = 6.689(3) Å, c = 17.403(6) Å, V = 674.4(5) Å3 , and Z = 3. The reﬂection
conditions (00l: l = 3n), together with the observed Rint in the diﬀerent Laue classes, point unequivocallyto the choice of the space group P32 21. The structure solution was then carried out in this space group.
The positions of most of the atoms (all the V positions and most of the S atoms) were determined bymeans of direct methods [29]. A least-squares reﬁnement on F2 using these heavy-atom positions andisotropic temperature factors produced an R factor of 0.123. Three-dimensional diﬀerence Fourier
synthesis yielded the position of the remaining sulfur atoms. The program Shelxl-97 [30] was used forthe reﬁnement of the structure. The site occupancy factor (s.o.f.) at the cation sites was allowed to vary
(V vs. structural vacancy) using scattering curves for neutral atoms taken from the International Tables
for Crystallography [30]. Four V sites (i.e., V1, V2, V3, and V4) were found to be partially occupied byvanadium (75%), while V5 and V6 were found to be fully occupied by V and ﬁxed accordingly. Sulfur
atoms were found on four fully occupied general positions leading to an ideal formula V7 S8 . Given thealmost identical partial occupancy of four V sites (i.e., 75%), we carefully checked either the possible
presence of twinning or the acentricity of the model. The lack of the inversion center was conﬁrmed
using the Flack parameter in Shelxl (0.07(1)) and the trigonal model was double-checked with the Platon
addsymm routine. At the last stage, with anisotropic atomic displacement parameters for all atoms andno constraints, the residual value settled at R1 = 0.0363 for 398 observed reﬂections (Fo > 4σ(Fo)) and
79 parameters and at R1 = 0.0537 for all 1289 independent reﬂections. Reﬁned atomic coordinates andisotropic displacement parameters are given in Table 3, whereas selected bond distances are reported
in Table 4. Crystallographic Information File (CIF) is deposited.
The calculated X-ray powder diﬀraction pattern (Table 5) was computed on the basis of theunit-cell data above and with the atom coordinates and occupancies reported in Table 3.

10Minerals 2020, 10, 245

Table 3. Atoms, site occupancy, fractional atom coordinates (Å), and isotropic atomic displacement
parameters (Å2 ) for eliopoulosite.
Atom

Site Occupancy

x/a

y/b

z/c

U iso

V1
V2
V3
V4
V5
V6
S1
S2
S3
S4

V0.749(6)V0.762(6)V0.755(6)V0.749(7)
V1.00
V1.00
S1.00
S1.00
S1.00
S1.00

0.0001(6)0.5005(6)0.0006(7)0.5010(7)0.4996(4)0.4997(5)0.1663(7)0.1659(7)0.1661(7)0.3333(6)00000.4997(5)0.4996(5)0.3343(6)0.3328(7)0.3331(6)0.6673(7)

2/3
2/3
1/6
1/6
0.16664(5)0.33327(5)0.08309(11)0.41639(11)0.74976(11)0.25009(11)0.0088(6)0.0095(6)0.0113(6)0.0187(8)0.0060(3)0.0120(4)0.0227(4)0.0226(4)0.0226(4)0.0228(4)Table 4. Bond distances (Å) in the structure of eliopoulosite.
V1-S2
V1-S3
V1-S1
mean

2.409(5) (×2)
2.411(3) (×2)
2.418(4) (×2)
2.413

V2-S1
V2-S2
V2-S4
mean

2.412(4) (×2)
2.413(5) (×2)
2.414(3) (×2)
2.413

V3-S3
V3-S2
V3-S1
mean

2.414(4) (×2)
2.418(3) (×2)
2.422(3) (×2)
2.418

V4-S4
V4-S2
V4-S3
mean

2.413(3) (×2)
2.417(4) (×2)
2.421(5) (×2)
2.417

V1-V3
V2-V5
V4-V6
V5-V6

2.9006(10) (×2)
2.9001(13) (×2)
2.9018(14) (×2)
2.8998(12) (×2)

V5-S1
V5-S4
V5-S1
V5-S2
V5-S4
V5-S3
mean

2.413(4)2.416(3)2.417(4)2.418(4)2.419(3)2.419(4)
2.417

V6-S1
V6-S4
V6-S3
V6-S3
V6-S2
V6-S4
mean

2.407(4)2.411(4)2.413(4)2.413(4)2.415(4)2.417(4)2.413

Table 5. Calculated X-ray powder diﬀraction data (d in Å) for eliopoulosite. The strongest observed
reﬂections are given in bold.
hkl

dcalc

Icalc

200
023
026
220
029
012
400
046
2212
246
600

2.89642.59142.04951.6723
1.60821.4503
1.4482
1.29571.09561.02420.9655

29
45
100
40
10
8
3
20
15
12
4

11

Minerals 2020, 10, 245

The chemical data (Table 1), yielding to a mean composition (wt.%) of S 41.78, V 54.11, Ni
1.71, Fe 1.1, Co 0.67, Mo 0.66, and total 100.03, were then normalized taking into account thestructural results. The empirical formula of eliopoulosite on the basis of Σatoms = 15 apfu is(V6.55 Ni0.19 Fe0.12 Co0.07 Mo0.04 )Σ = 6.97 S8.03 . The simpliﬁed formula is (V,Ni,Fe)7 S8 and the ideal formulais V7 S8 , which corresponds to V 58.16, S 41.84, and total 100 (wt.%).
The structure can be considered as a twelve-fold superstructure (2a × 2a × 3c) of the NiAs-type
subcell with V atoms octahedrally coordinated by S atoms and sulfur located in a trigonal prism.
There are six octahedral layers in the structure with rods of fully occupied V sites alternate to rodsof partially occupied sites, so that every layer contains 3.5 vanadium atoms. In successive layers
rods are directed along [110], [100], and [010] accordingly to the threefold screw axis. As shown in
Figure 5, the distribution of fully (blue in color) and partially (pale blue) occupied sites determine thesuperstructure along the a1 and a2 axes. Diﬀerences in electron density between fully and partiallyoccupied sites are indeed rather modest (22 vs. 16.5): Accordingly, hkl reﬂections with h,k = 2n + 1 arerather weak (<I/σ(I)> = 27.1 vs. 4890.3).
Since the octahedral vacancies’ distribution within the layers at z = 0, z = 1/6, and z = 1/3 isperfectly replicated at layers at z = 1/2, z = 2/3, and z = 5/6 (Figure 5), no contribution to hkl reﬂections
with l = 2n + 1 is given by metal atoms. However, their intensities (<I/σ(I)> = 564.3 vs. 3349.2) are evenhigher than those leading to doubling of a-axis, due to the diﬀerent position of sulfur atoms related tothe diﬀerent octahedral orientation in the layer at z = 0 with respect to z = 1/2 (Figure 4).

Figure 4. Octahedral layers in eliopoulosite projected down [001]. Layer at z = 0 (top of picture) repeatsat z = 1/3 and z = 2/3; layer at z = 1/2 (bottom of picture) repeats at z = 1/6 and z = 5/6. Blue and pale
blue colors represent full and partial occupancy. Sulfur atoms are depicted in yellow. The unit cell andthe orientation of the structure are indicated.

12

Minerals 2020, 10, 245

The eliopoulosite structure shows strong analogies with the superstructures observed in thepyrrhotite-group of minerals. It is identical to the structure inferred for synthetic VS1.125 [31]. The mean
<V-S> bond distances in eliopoulosite (in the range 2.413–2.418 Å) are in excellent agreement with thosefound for synthetic vanadium sulﬁdes. To achieve the charge balance, one should hypothesize thepresence of both divalent and trivalent vanadium in eliopoulosite (i.e., V2+ 5 V3+ 2 S8 ). The octahedral sites,
however, do not show any signiﬁcant diﬀerence symptomatic of a V2+ -V3+ ordering. Likewise, we did
not ﬁnd any evidence of ordering of the minor substituents (i.e., Ni and Fe) at a particular structuralsite. Furthermore, eliopoulosite shows, for the analogy of stoichiometry, unit cell and space group,
close relationships with the metastable, trigonal form of 3C-Fe7 S8 (a = 6.852(6), c = 17.046(2), P31 21) [32].
However, the distribution of vacancies is quite diﬀerent (Figure 5b): As in other types of pyrrhotites [33],in 3C-Fe7 S8 octahedral vacancies are completely ordered on one position alone so that the layers
containing a void (i.e., three metal atoms) are alternating with fully occupied layers (i.e., four
metal atoms).

Figure 5. Metal distribution in eliopoulosite (a) and in synthetic 3C-Fe7 S8 pyrrhotite (b). Blue and pale
blue circles refer to metal positions with full and partial (75%) occupancy, respectively. White circles
represent completely empty sites in synthetic 3C-Fe7 S8 pyrrhotite [32]. The unit cell and the orientationof the structure are indicated.

13

Minerals 2020, 10, 245

Eliopoulosite does not correspond to any valid or invalid unnamed mineral [34]. According toliterature data [1–8] eliopoulosite is the ninth mineral containing V and S that has been accepted by IMA.
Noteworthy, eliopoulosite is the second V2+ -bearing mineral, the ﬁrst being dellagiustaite, Al2 VO4 [35].
The presence of divalent V requires extremely reducing conditions (well below the iron-wüstite buﬀer).
Furthermore, based on its chemical composition, eliopoulosite is the second sulﬁde that contains only
V as a major component discovered so far. Recently, Ivanova et al. [36] reported on the presenceof a V,Cr,Fe-sulﬁde in the extremely reduced assemblage of a rare CB chondrite. Due to its small
size (one micron) the grain was studied only by EBSD (Electron BackScattered Diﬀraction) and nocrystallographic or optical data were provided. However, its ideal composition (V,Cr,Fe)4 S5 seems
diﬀerent compared to that of eliopoulosite in terms of S content and V/Fe ratio and the presence ofabundant Cr.
Interestingly, Selezneva et al. [37] recently showed that synthetic V7 S8 shows diﬀerences in
the magnetic behavior with respect to trigonal Fe7 S8 . Indeed, V7 S8 is observed to exhibit aPauli-paramagnetic behavior opposite to the classic ferrimagnetic ordering usually observed inpyrrhotite. Unfortunately, the small amount of the natural material precludes any possible measurementof both the magnetic susceptibility and magnetization. Such an experiment would have allowed us toverify if the amount of structural vacancies observed for eliopoulosite are enough to compensate themagnetic moments to avoid the ferromagnetic order.
6. Remarks on the Origin of Eliopoulosite
Eliopoulosite was found in the same sample in which tsikourasite and grammatikopoulosite werediscovered, and the three minerals occur in the same mineralogical assemblage [9,10]. Therefore, we
can argue that they formed under the same chemical-physical condition (i.e., in a strongly reducing
environment). This assumption is fully supported by the chemical composition of eliopoulosite,that points to a low valence state for V. However, as already proposed for the origin of tsikourasite
and grammatikopoulosite, it is still not possible to provide a conclusive model to explain exhaustively
the genesis of eliopoulosite. The following hypothesis can be formulated, since all of them imply thepresence of a reducing environment [9,10,38–43]: i) low-temperature alteration of chromitite duringserpentinization process; ii) high-temperature reaction of the chromitites with reducing ﬂuids, at mantle
depth; iii) post-orogenic surface lightning strike; or iv) meteorite impact.
However, the probability of intercepting a fragment of a meteorite or a fulgurite in the Otrhys
ophiolite during the sampling of the studied chromitite seems very unlikely, although a V-bearing
sulﬁde has been reported in a meteorite [36].
Supplementary Materials: The following are available online at http://www.mdpi.com/2075-163X/10/3/245/s1,
CIF: eliopoulosite.
Author Contributions: L.B., F.Z., and P.B. wrote the manuscript; F.Z. performed the chemical analyses; L.B. and
P.B. performed the diﬀraction experiments; T.G. and B.T. provided the concentrate sample and information on thesample provenance and petrography of Othrys chromitite; G.G. discussed the chemical data and C.S. obtained theoptical data. All the authors provided support in the data interpretation and revised the manuscript. All authors
have read and agreed to the published version of the manuscript.
Funding: The authors are grateful to the University Centrum for Applied Geosciences (UCAG) for the access to
the E. F. Stumpﬂ electron microprobe laboratory. SGS Mineral Services, Canada, is thanked to have performed theconcentrate sample. S. Karipi is thanked for sample collection and participating in the ﬁeld work. L.B. thanks
MIUR, project “TEOREM deciphering geological processes using Terrestrial and Extraterrestrial ORE Minerals”,
prot. 2017AK8C32 (PI: Luca Bindi).
Acknowledgments: The authors acknowledge Ritsuro Miyawaki, Chairman of the CNMNC and its members forhelpful comments on the submitted new mineral proposal. Many thanks are due to the editorial staﬀ of Minerals.
Conﬂicts of Interest: The authors declare no conﬂicts of interest.

14

Minerals 2020, 10, 245

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Petrascheck, W., Karamata, S., Eds.; Theophrastus Publ. S.A.: Athens, Greece, 1986; pp. 129–159.
Ifandi, E.; Zaccarini, F.; Tsikouras, B.; Grammatikopoulos, T.; Garuti, G.; Karipi, S.; Hatzipanagiotou, K. First
occurrences of Ni-V-Co phosphides in chromitite of Agios Stefanos mine, Othrys ophiolite, Greece. Oﬁoliti
2018, 43, 131–145.
Zaccarini, F.; Ifandi, E.; Tsikouras, B.; Grammatikopoulos, T.; Garuti, G.; Mauro, D.; Bindi, L.; Stanley, C.
Occurrences of new phosphides and sulﬁde of Ni, Co, V, and Mo from chromitite of the Othrys ophiolitecomplex (Central Greece). Per. Mineral. 2019, 88. [CrossRef]
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x-ray-diffraction-and-elemental-analysis/single-crystal-x-ray-diffraction/sc-xrd-software/apex3.html (accessedon 6 March 2020).
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products/x-ray-diffraction-and-elemental-analysis/single-crystal-x-ray-diffraction/sc-xrd-software/apex3.html(accessed on 6 March 2020).
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Union of Crystallography: Chester, UK, 1992; Volume 3.
Grønvold, F.; Haraldsen, H.; Pedersen, B.; Tufte, T. X-ray and magnetic study of vanadium sulﬁdes in therange V5 S4 to V5 S8 . Rev. Chim. Minéral. 1969, 6, 215.
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Unnamed Minerals of the IMA Commission on New Minerals, Nomenclature and Classiﬁcation. Can. Mineral.
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Ivanova, M.A.; Ma, C.; Lorenz, C.A.; Franchi, I.A.; Kononkova, N.N. A new unusual bencubbinite (cba),
Sierra Gorda 013 with unique V-rich sulﬁdes. Met. Plan. Sci. 2019, 54, 6149.
Selezneva, N.V.; Ibrahim, P.N.G.; Toporova, N.M.; Sherokalova, E.M.; Baranov, N.V. Crystal structure andmagnetic properties of pyrrhotite-type compounds Fe7-y Vy S8 . Acta Phys. Polon. 2018, A133, 450–452.
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Etiope, G.; Tsikouras, B.; Kordella, S.; Ifandi, E.; Christodoulou, D.; Papatheodorou, G. Methane ﬂux andorigin in the Othrys ophiolite hyperalkaline springs, Greece. Chem. Geol. 2013, 347, 161–174. [CrossRef]
Etiope, G.; Ifandi, E.; Nazzari, M.; Procesi, M.; Tsikouras, B.; Ventura, G.; Steele, A.; Tardini, R.; Szatmari, P.
Widespread abiotic methane in chromitites. Sci. Rep. 2018, 8, 8728. [CrossRef]
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minerals
ArticleGrammatikopoulosite, NiVP, a New Phosphide fromthe Chromitite of the Othrys Ophiolite, Greece
Luca Bindi 1 , Federica Zaccarini 2, *, Elena Ifandi 3,4 , Basilios Tsikouras 3 , Chris Stanley 5 ,Giorgio Garuti 2 and Daniela Mauro 6
1
2
3
4
5
6

*

Dipartimento di Scienze della Terra, Università degli Studi di Firenze, I-50121 Florence, Italy;luca.bindi@uniﬁ.it
Department of Applied Geological Sciences and Geophysics, University of Leoben, A-8700 Leoben, Austria;
giorgio.garuti1945@gmail.com
Faculty of Science, Physical and Geological Sciences, Universiti Brunei Darussalam, BE 1410 Gadong,
Brunei Darussalam; selena.21@windowslive.com (E.I.); basilios.tsikouras@ubd.edu.bn (B.T.)
Department of Geology, Section of Earth Materials, University of Patras, 265 00 Patras, Greece
Department of Earth Sciences, Natural History Museum, London SW7 5BD, UK; c.stanley@nhm.ac.uk
Dipartimento di Scienze della Terra, Università degli Studi di Pisa, I-56126 Pisa, Italy;
daniela.mauro@dst.unipi.it
Correspondence: federica.zaccarini@unileoben.ac.at; Tel.: +43-(0)3842-402-6218

Received: 22 December 2019; Accepted: 27 January 2020; Published: 31 January 2020

Abstract: Grammatikopoulosite, NiVP, is a new phosphide discovered in the podiform chromitite andhosted in the mantle sequence of the Othrys ophiolite complex, central Greece. The studied samples
were collected from the abandoned chromium mine of Agios Stefanos. Grammatikopoulosite forms
small crystals (from 5 μm up to about 80 μm) and occurs as isolated grains. It is associatedwith nickelphosphide, awaruite, tsikourasite, and an undetermined V-sulphide. It is brittle
and has a metallic luster. In plane-polarized light, it is creamy-yellow, weakly bireﬂectant,with measurable but not discernible pleochroism and slight anisotropy with indeterminate rotation
tints. Internal reﬂections were not observed. Reﬂectance values of mineral in air (R1 , R2 in %)are: 48.8–50.30 at 470 nm, 50.5–53.5 at 546 nm, 51.7–55.2 at 589 nm, and 53.2–57.1 at 650 nm.
Five spot analyses of grammatikopoulosite give the average composition: P 19.90, S 0.41, Ni 21.81,
V 20.85, Co 16.46, Mo 16.39, Fe 3.83, and Si 0.14, total 99.79 wt %. The empirical formula ofgrammatikopoulosite—based on Σ(V + Ni + Co + Mo + Fe + Si) = 2 apfu, and taking into account thestructural results—is (Ni0.57 Co0.32 Fe0.11 )Σ1.00 (V0.63 Mo0.26 Co0.11 )Σ1.00 (P0.98 S0.02 )Σ1.00 . The simpliﬁed
formula is (Ni,Co)(V,Mo)P and the ideal formula is NiVP, which corresponds to Ni 41.74%, V 36.23%,
P 22.03%, total 100 wt %. The density, calculated on the basis of the empirical formula and single-crystal
data, is 7.085 g/cm3 . The mineral is orthorhombic, space group Pnma, with a = 5.8893(8), b = 3.5723(4),c = 6.8146(9) Å, V = 143.37(3) Å3 , and Z = 4. The mineral and its name have been approved by theCommission of New Minerals, Nomenclature and Classiﬁcation of the International Mineralogical
Association (IMA 2019-090). The mineral honors Tassos Grammatikopoulos, geoscientist at the SGS
Canada Inc., for his contribution to the economic mineralogy and mineral deposits of Greece.
Keywords: grammatikopoulosite; phosphide; chromitite; Agios Stefanos mine; Othrys;
ophiolite; Greece

1. Introduction
Natural phosphides are very rare phases, representing only 3% of the minerals approvedby the International Mineralogical Association (IMA) ([1] and references therein). Most of thesenatural phosphides have been discovered and described in meteorites. Only recently have several new

Minerals 2020, 10, 131; doi:10.3390/min10020131

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Minerals 2020, 10, 131

phosphides been discovered in terrestrial rocks, most of which are associated with the pyrometamorphic
rocks of the Hatrurim Formation, Southern Levant [2]. Phosphides are documented in podiform
chromitites hosted in the mantle sequence of ophiolite complexes of Greece and Russia [3–6] but
are very rare. The phosphide found in the Russian chromitite was found in situ, in contact withserpentine of the altered interstitial silicates. The phosphides associated with the Greek chromitite were
documented in heavy concentrates. Recently, the new mineral tsikourasite (Mo3 Ni2 P1 + x (x < 0.25)
was discovered in heavy mineral concentrates from a chromitite from the Othrys ophiolite (Greece)) [1].
Further investigation of the same mineral concentrates led to the discovery of a second new mineral.
Quantitative chemical analysis and crystal structure proved that the studied phase is a new phosphide,characterized by the simpliﬁed formula (Ni,Co)(V,Mo)P and by the ideal formula NiVP, whichcorresponds to Ni 41.74%, V 36.23%, P 22.03%, total 100 wt %. The mineral is orthorhombic, spacegroup Pnma, with a = 5.8893(8), b = 3.5723(4), c = 6.8146(9) Å, and Z = 4. The mineral and its namehave been approved by the Commission of New Minerals, Nomenclature and Classiﬁcation of theInternational Mineralogical Association (IMA 2019-090). The mineral honors Tassos Grammatikopoulos(b. 1966), a geoscientist at the SGS Canada Inc., for his contribution to the investigation of economic
mineralogy and mineral deposits of Greece. Holotype material is deposited in the MineralogicalCollection of the Museo di Storia Naturale, Università di Pisa, Via Roma 79, Calci (Pisa, Italy), undercatalogue number 19,911.
2. Geological Background and Occurrence of Grammatikopoulosite
Grammatikopoulosite was discovered in a heavy mineral concentrate, which was prepared fromchromitite specimens hosted in the mantle sequence of the Mesozoic Othrys ophiolite, central Greece(Figure 1A–C).
Othrys ophiolite is a complete but dismembered suite (Mirna Group) and consists of three structuralunits: the uppermost succession with variably serpentinized peridotites, which is structurally boundedby an ophiolite mélange; the intermediate Kournovon dolerite, including cumulate gabbro and local
rhyolite; and the lower Sipetorrema Pillow Lava unit including also basaltic ﬂows, siltstones, and chert.
The Mirna Group constituted multiple inverted thrust sheets [7], which were eventually obducted onto
the Pelagonian Zone during the Late Jurassic–Early Cretaceous [8–14]. Three types of basalts with
diﬀerent geochemical signatures have been described: (i) alkaline within-plate (WPB), (ii) normal-type
mid-ocean ridge (N-MORB), and (iii) low-K tholeiite (L-KT). A biostratigraphic investigation indicatedthat radiolarites associated with N-MORB were deposited in the Middle and Late Triassic. Radiolarites
deposited over the L-KT basalts are Early Carnian–Middle Norian–Late Norian in age [15]. N-MORB
erupted during the Middle–Late Triassic period. The L-KT basalts erupted during the Middle–Late
Triassic period, in part simultaneously with N-MORB, in the ocean–continent transition zone, close
to the rifted continental margin. Finally, the alkaline WPB are interpreted to have formed in oceanic
seamounts or in the ocean–continent transition zone adjacent to the rifted continental margin [15].
Electron microprobe analyses of the assemblages of these chromitites revealed that the studied
spinel-supergroup minerals are magnesiochromites [1,5,6]. However, spinel mineral chemistry is rather
heterogeneous with Cr2 O3 (44.96–51.64 wt %), Al2 O3 (14.18–20.78 wt %), MgO (13.34–16.84 wt %),and FeO (8.3–13.31 wt %). The calculated Fe2 O3 ranges from 6.72 to 9.26 wt %. The amounts of MnO(0.33–0.60 wt %), V2 O3 (0.04–0.30 wt %), ZnO (up to 0.07 wt %), and NiO (0.03–0.24 wt %) exhibit minor
variations. The TiO2 content is low (0.03–0.23 wt %), in agreement with the typical value reported forthe mantle-hosted podiform chromites. The Cr/(Cr + Al) ratios of the investigated magnesiochromite
are lower than those of chromites formed in the supra-subduction zone (SSZ) and those related toboninitic melt [1,5,6].

20

Minerals 2020, 10, 131

Figure 1. Location of the Othrys complex in Greece: (A) general geological map of the Othrys ophiolite
showing the location of the Agios Stefanos chromium mine; (B) and (C) detailed geological setting of
the Agios Stefanos area (modiﬁed after [5,10,16,17]).

3. Analytical Methods
The processing and recovery of the heavy minerals were carried out by treating about 10 kg of
massive chromitite at SGS Mineral Services, Canada, following the procedure described by several
authors [1,5,6,18]. Approximately 10 kg of chromitite were collected in the abandoned mine of Agios
Stefanos. Subsequently, they were crushed and blended to generate a homogenous composite sample.
About 500 g of the composite sample were riﬄed and stage crushed to a P80 (80% passing) of 75 μm.
After, the sample was treated with heavy liquid (the density of the heavy liquids was 3.1 g/cm3 )to separate the heavy minerals. This procedure produced a heavy and light mineral concentrate.
The heavy fraction was selected and processed with a superpanner. This method is designed forsmall amounts of sample and is closely controlled, leading to very eﬀective separation. It consists of atapering triangular deck with a “V” shape cross section. The table reproduces the concentrating actionof a gold pan. Initially, the sample is swirled to stratify the minerals. Then, the heaviest minerals
settle to the bottom and are deposited on the deck surface. The less dense material moves towards thetop, overlying the heavy minerals. The operation of the deck is then changed to a rapid reciprocal
motion, with an appropriate “end-knock” at the up-slope end of the board, and a steady ﬂow ofwash water is introduced. The “end-knock” forces the heavy minerals to migrate to the up-slope
end of the deck. The wash water carries the light minerals to move to the narrower, down-slope
end of the deck. The heaviest fractions were split into the heaviest fraction (tip) followed by a lessdense fraction (middling). The “tip” and the “middlings” of the superpanner are the densest fractions
and included liberated grains of chromite, sulphides, alloys, and phosphides, while the lighter tailconsisted of a particle mixture of chromite and silicates. No source of contamination is likely during
21

Minerals 2020, 10, 131

sample collection and subsequent treatment. The heavy minerals were prepared in epoxy blocks,
and then polished for mineralogical examination. Quantitative chemical analyses and acquisitionof back-scattered electron images of grammatikopoulosite were performed with a JEOL JXA-8200
electron microprobe, installed in the E. F. Stumpﬂ laboratory, Leoben University, Austria, operating in
Wavelength Dispersive Spectrometry (WDS) mode. Major and minor elements were determined at
a 20 kV accelerating voltage and a 10 nA beam current, with 20 s as the counting time for the peakand 10 s for the backgrounds. The beam diameter was about 1 μm in size. For the WDS analyses,
the following lines and diﬀracting crystals were used: P, S, Si = (Kα, PETJ), V, Fe, Co, Ni = (Kα, LIFH),
and Mo = (Lα, PETJ). The following standards were selected: synthetic Ni3 P for Ni and P, molybdenite
for S and Mo, synthetic Fe3 P for Fe, synthetic metallic vanadium for V, skutterudite for Co, quartz
for Si, and chromite for Cr. The ZAF correction method was applied. Automatic corrections were
performed for interferences P-Mo, Cr-V, and Mo-S. Representative analyses of grammatikopoulosite
are listed in Table 1.
Table 1. Chemical data (wt % of elements) for grammatikopoulosite.
Sample

P

S

Ni

V

Co

Mo

Fe

Si

Total

VP40-1
VP40-2
VP40-3
VP40-4
VP40-5average

20.38
19.83
19.65
19.65
20.01
19.90

0.41
0.42
0.39
0.40
0.41
0.41

21.98
21.70
21.72
21.95
21.69
21.81

21.02
20.48
20.73
21.05
20.98
20.85

16.33
16.66
16.51
16.37
16.45
16.46

16.72
16.36
16.35
16.31
16.20
16.39

3.82
3.83
3.86
3.85
3.78
3.83

0.14
0.14
0.13
0.13
0.16
0.14

100.79
99.41
99.33
99.71
99.67
99.79

Single-crystal and powder X-ray diﬀraction data were collected at the University of Florence
using a Bruker D8 Venture equipped with a Photon II CCD detector, with graphite-monochromatized
MoKα radiation (λ = 0.71073 Å) using a crystal fragment hand-picked from the polished section undera reﬂected light microscope (cif ﬁle, see Supplementary Materials). The crystal (about 80 μm in size)
was carefully and repeatedly washed in acetone. It did not show any other visible phase attached tothe surface. Single-crystal X-ray diﬀraction intensity data were integrated and corrected for standard
Lorentz polarization factors with the software package Apex3 [19,20].
The reﬂectance measurements on grammatikopoulosite were carried out using a WTiC standard
and a J&M TIDAS diode array spectrophotometer at the Natural History Museum of London, UK.
4. Physical and Optical Properties
More than 30 grains of grammatikopoulosite were found in the studied polished sections.
Grammatikopoulosite occurs as anhedral to subhedral grains. Most of them were less than 10 μm insize and only two of them were up to about 80 μm. Grammatikopoulosite consists of single (Figure 2A)or poly-phase grains associated with other minerals, such as tsikourasite, nickelphosphide, awaruite,
and a V-sulphide, which likely represents another new mineral (Figure 2B,C). Compositions of theseminerals have been reported in previous papers [5,6]. In plane-polarized light, grammatikopoulosite
is creamy-yellow, weakly bireﬂectant, with measurable but not discernible pleochroism and slightanisotropy with indeterminate rotation tints. Internal reﬂections were not observed. Reﬂectance valuesof the mineral in air (R in %) are reported in Table 2 and in Figure 3. Density was not measured
because of the small amount of available material and the presence of ﬁne intergrowths with awaruite.
The calculated density is equal to 7.085 g·cm−3 , based on the empirical composition and unit–cellvolume reﬁned from single-crystal XRD data.

22

Minerals 2020, 10, 131

Figure 2. Digital image in reﬂected plane-polarized light (A,B) and a back-scattered electron image(C) showing grammatikopoulosite from the chromitite of Agios Stefanos. Abbreviations: Grm =grammatikopoulosite, Tsk = tsikourasite, VS = V-sulphide, Aw = awaruite, Npd = nickelphosphide,
Epx = epoxy. Scale bar = 50 microns.

Figure 3. Reﬂectance data for grammatikopoulosite.

23

Minerals 2020, 10, 131

Table 2. Reﬂectance values for grammatikopoulosite. The values required by the Commission on Ore
Mineralogy are given in bold.
λ nm

R1

λ nm

R2

400
420
440
460
470
480
500
520
540
546
560
580
589
600
620
640
650
680
700

47.6
47.9
48.3
48.6
48.8
49.0
49.4
49.9
50.3
50.5
50.9
51.4
51.7
51.9
52.4
53.0
53.2
53.8
54.2

400
420
440
470
470
480
500
520
540
546
560
580
589
600
620
640
650
680
700

48.8
49.1
49.4
49.9
50.3
50.7
51.5
52.4
53.3
53.5
54.1
54.9
55.2
55.5
56.2
56.8
57.1
58.0
58.6

5. Chemical Composition and X-Ray Crystallography
Chemical composition and X-ray data reveal that the empirical formula of grammatikopoulosite,
based on Σ(V + Ni + Co + Mo + Fe + Si) = 2 apfu and taking into account the structural results(see below), is (Ni0.57 Co0.32 Fe0.11 )Σ1.00 (V0.63 Mo0.26 Co0.11 )Σ1.00 (P0.98 S0.02 )Σ1.00 . The simpliﬁed formula is(Ni,Co)(V,Mo)P and the ideal formula is NiVP, which corresponds to the composition of Ni = 41.74%,
V = 36.23%, and P 22.03% (total 100 wt %).
A small grammatikopoulosite grain (about 80 μm in size) was handpicked from the polished section
under a reﬂected light microscope and mounted on a 5 μm diameter carbon ﬁber, which was, in turn,attached to a glass rod in preparation for the single-crystal X-ray diﬀraction study. The fragment consistsof crystalline grammatikopoulosite associated with minor, ﬁne-grained polycrystalline awaruite. A totalof 493 unique reﬂections were collected up to 2θ = 80.33◦ . The mineral is orthorhombic, space groupPnma, with a = 5.8893(8), b = 3.5723(4), c = 6.8146(9) Å, and Z = 4. Given the similarity in unit–cellvalues and space groups, the structure was reﬁned starting from the atomic coordinates reported
for allabogdanite [21] using the software Shelxl-97 [22]. The site occupancy factor (s.o.f.) at the twocation sites was allowed to vary (Ni vs. Mo) using scattering curves for neutral atoms taken fromthe International Tables for Crystallography [23], leading to 27.5 and 28.4 e− for M1 and M2 sites,
respectively. The P site was found to be fully occupied (P vs. structural vacancy) by phosphorous(site scattering = 15.0 e− ) and ﬁxed accordingly. Taking into account the site distribution observedin ﬂorenskyite [24], allabogdanite [21], and andreyivanovite [25], and in most Co2 Si-structure-type
synthetic compounds [26], we assigned Ni, Fe, and Co to ﬁll the M1 site (i.e., Ni0.57 Co0.32 Fe0.11 ) and all
the other elements (i.e., V0.63 Mo0.26 Co0.11 ) to the M2 site. The mean electron numbers calculated withsuch a site distribution were identical (27.36 and 28.38 e− ) to those obtained in the reﬁnement. For thisreason, the subsequent cycles of reﬁnement were run with the above constrained site populations,yielding an R1 = 0.0276 for 465 reﬂections with Fo > 4σ(Fo) and R1 = 0.0291 for all the 493 independent
reﬂections and 19 parameters. Reﬁned atomic coordinates and isotropic displacement parameters are
given in Table 3, whereas selected bond distances are reported in Table 4.
X-ray powder diﬀraction data for grammatikopoulosite (Table 5) were obtained with a Bruker D8
Venture equipped with a Photon II CCD detector, with graphite-monochromatized CuKα radiation(λ = 1.54138 Å). The least squares reﬁnement gave the following values: a = 5.8088(2), b = 3.5993(2),c = 6.8221(3) Å, and V = 142.634(8) Å3 . The calculated powder diﬀraction pattern obtained using thesite occupancies and atomic coordinates is reported in Table 3.

24

Minerals 2020, 10, 131

Table 3. Atoms, site occupancies, atom coordinates, and isotropic displacement parameters (Å2 )for grammatikopoulosite.
Atom

Site Occupancy

x/a

y/b

z/c

U iso

M1
M2
P

Ni0.57 Co0.32 Fe0.11
V0.63 Mo0.26 Co0.11
P1.00

0.35709(6)0.47087(6)0.23639(12)

1
4
1
4
1
4

0.93703(5)0.33109(5)0.62449(10)0.00578(10)0.00595(9)0.00547(13)Table 4. Selected bond distances (Å) for grammatikopoulosite.
Atoms

Bond Distance

M1–P
M1–P (×2)
M1–P
M1–M1 (×2)
M1–M2 (×2)
M1–M2 (×2)
M1–M2
M1–M2

2.2453(8)2.2639(5)2.2728(8)2.6000(6)2.7280(5)2.7487(5)2.7677(5)2.7696(5)

M2–P
M2–P (×2)
M2–P (×2)
M2–M1 (×2)
M2–M1 (×2)
M2–M1
M2–M1
M2–M2 (×2)

2.4299(8)2.5008(6)2.5811(6)2.7280(5)2.7487(5)2.7677(5)2.7696(5)2.9339(6)Table 5. Observed and calculated X-ray powder diﬀraction data (d in Å) for grammatikopoulosite.
The strongest observed reﬂections are given in bold.
Indices

1

2

hkl

dobs

Iobs

dcalc

Icalc

101
002
102
111
201
112
210
201
211
103
013
301
113
020
004
302
213
114
303
400
123
322
314
215
511
513

4.43
2.950
2.785
2.699
2.273
2.269
2.230
2.157
2.118
1.915
1.888
1.824
1.784
1.702
1.700
1.608
1.489
1.482
1.470
1.367
1.233
1.211
1.170
1.102
1.005

10
20
25
5
60
10
10
100
25
15
10
15
20
10
15
10
5
5
5
10
10
5
10
5
5

4.45593.4073
2.9493
2.78722.70312.27432.27222.22792.1555
2.1194
1.91681.8864
1.8227
1.78611.70361.70101.60651.48781.48531.47231.36581.23181.2106
1.16881.10381.0035

14
5
19
24
5
65
8
9
100
27
14
12
20
21
10
22
7
6
6
6
7
9
6
9
6
5

1 = observed diﬀraction pattern; 2 = calculated diﬀraction pattern obtained with the atomic coordinates reported inTable 3 (only reﬂections with Irel ≥ 4 are listed).

25

Minerals 2020, 10, 131

6. Description of the Structure and Relations to Other Species
In the structure of grammatikopoulosite (Figure 4), M1 links four P atoms and eight M2(Figure 5—left), whereas M2 links ﬁve P, six M1, and two M2 (Figure 5—right). The metal–phosphorous
distances are much shorter in the M1 coordination sphere than in that of M2 (Table 4). Interestingly,
if only the M–P distances are considered in the coordination polyhedra of the M atoms, M1P4 tetrahedra(Figure 6) forming corner-sharing chains along the b-axis or M2P5 square pyramids forming zig-zag
chains along the a-axis (Figure 7) can be observed.

Figure 4. The crystal structure of grammatikopoulosite projected down [10]. Blue, green, and red
circles refer to M1, M2, and P, respectively.

Figure 5. Coordination environment of M1 (left) and M2 (right) sites in the crystal structure ofgrammatikopoulosite. Colors and symbols as in Figure 4.

Figure 6.
Polyhedral representation of the M1P4 tetrahedra in the crystal structure ofgrammatikopoulosite. Colors and symbols as in Figure 4.

26

Minerals 2020, 10, 131

Figure 7. Polyhedral representation of the M2P5 distorted pyramids in the crystal structure ofgrammatikopoulosite. Colors and symbols as in Figure 4.

It was reported that the departure from the hexagonal, high-temperature P-62m structure(barringerite structure) was linked to some ordering between the metals at the M sites. This was not
immediately evident in both allabogdanite [21] and andreyivanovite [25] because of the close site
scattering of Fe, Ni, and Cr. On the contrary, the ordering is evident if we take into consideration
ﬂorenskyite [24] and grammatikopoulosite with ideal formulae FeTiP and NiVP, respectively. Ti and V
preferentially enter the larger M2 site in agreement with their larger metallic radius (1.47 and 1.35 Å,
respectively [27]).
Grammatikopoulosite does not correspond to any valid or invalid unnamed mineral [28] andit belongs to the group of natural phosphides with the Co2 Si orthorhombic Pnma structure, i.e.,
ﬂorenskyite [24], allabogdanite [21], and andreyivanovite [25].
7. Discussion and Genetical Implications
Grammatikopoulosite, similar to tsikourasite, is natural in origin despite the fact that both phases
were found in heavy mineral concentrates and not in situ [1]. This assumption is supported by
the following observations: during the sampling, the only tool used to collect the chromitite wasa steel hammer; the concentrates were obtained in a laboratory that uses a grinder manufactured
by TM Engineering, with media Alloy 1 composed of Cr and Mo, and does not contain Ni, V,
Co, Fe, or P; the used material is very hard and is recognized to be best for applications whereabrasion and impact is the norm; and the chemical composition of grammatikopoulosite indicates
crystallization under reducing conditions, similar to what has been previously proposed for otherphosphides found in ophiolitic chromitites [1,3–6]. According to literature data, the crystallization
temperature of most of the terrestrial phosphides is generally comprised between 700 ◦ C and 1150 ◦ C [4].
However, a conclusive model to explain the origin of grammatikopoulosite cannot be provided yet,since it may have a terrestrial origin or represent fragments of a meteorite. The following models,
which imply the crystallization in a terrestrial environment from low to high temperatures, can besuggested for the origin of grammatikopoulosite: (i) alteration of chromitite during sub-oceanic oron-land serpentinization at temperatures between 400 ◦ C and 150 ◦ C, which is a typical reducing
process [29]; (ii) reaction of the chromitite with high-temperature reducing ﬂuids in the mantle;
and (iii) interaction of the chromitite and their host serpentinized peridotites with a surface lightning
strike. Phosphide minerals occurring in ophiolitic chromitites in Othrys and Gerakini–Ormylia(Greece), as well as in Alapaevsk (Russia), are closely associated with highly reducing phases such as
awaruite, typically forming during serpentinization [1,3–6]. Likewise, grammatikopoulosite is also
associated with awaruite and other reducing phases, likely suggesting a low-temperature origin duringserpentinization. This hypothesis is strongly supported by the discovery of a Ni-phosphide in situ in27

Minerals 2020, 10, 131

the serpentine-rich matrix of the Alapaevsk chromitite of Russia [3]. Recently, Etiope et al. [30,31] haveshown that the Agios Stefanos chromitite is the source rock of abiotic CH4 measured in the west Othrys
springs. These authors argued that CH4 formation took place at temperatures below 150 ◦ C via Sabatier
reaction during oceanic and on-land serpentinization. This observation provides further evidence that
reducing conditions can be achieved at a low temperature during the alteration of mantle-derived
rocks. Alternatively, Xiong et al. [32] suggested that reducing conditions can be achieved in mantle
rocks, including chromitites, in the shallow lithosphere due to the interaction of mantle-derived ﬂuids
enriched in CH4 and H2 with basaltic melt. However, no phosphides are reported in the ultra-reduced,
high-temperature mineralogical assemblage described by Xiong et al. [32]. Natural phosphides andother ultra-reduced minerals have been rarely described in few fulgurites [33,34] but they are very
common accessory phases in meteorites [21,24,25,35–39] or. The rare and localized occurrence of the
Othrys phosphides may provide some support to the two last models. However, the probability tohave collected a fragment of a meteorite or a fulgurite in the Othrys ophiolite during the sampling of
the studied chromitite seems improbable, and for the discrimination of the meteoritic materials, a moredetailed isotopic study may be needed. To conclude, an indisputable genetic model to explain how
grammatikopoulosite is crystallized still needs further investigation.
Supplementary Materials: The following are available online at http://www.mdpi.com/2075-163X/10/2/131/s1,
CIF: grammatikopoulosite.
Author Contributions: L.B. and F.Z. wrote the manuscript; F.Z. and L.B. performed the chemical analyses andthe diﬀraction experiments, respectively; E.I. and B.T. provided the concentrate sample and information on thesample provenance and petrography of Othrys chromitite; G.G. and D.M. discussed the chemical data; and C.S.
obtained the optical data. All authors provided support in the data interpretation and revised the manuscript.
All authors have read and agreed to the published version of the manuscript.
Funding: The authors are grateful to the University Centrum for Applied Geosciences (UCAG) for the access to the
E.F. Stumpﬂ electron microprobe laboratory. The authors also thank SGS Mineral Services, Canada, for performing
the concentrate sample. S. Karipi is thanked for collecting the sample and participating in the ﬁeld work. L.
Bindi thanks MIUR, project “TEOREM deciphering geological processes using Terrestrial and Extraterrestrial ORE
Minerals”, prot. 2017AK8C32 (PI: Luca Bindi).
Acknowledgments: The authors acknowledge Ritsuro Miyawaki, Chairman of the CNMNC and its members forhelpful comments on the submitted new mineral proposal. We are very grateful to the four anonymous referees
for constructive comments that improved the manuscript. Many thanks are due to the editorial staﬀ of Minerals
and to the consulted referees. This paper is dedicated to the memory of our friend and colleague Demetrios.
Conﬂicts of Interest: The authors declare no conﬂicts of interest.

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minerals
ArticlePodiform Chromitites and PGE Mineralization in the
Ulan-Sar’dag Ophiolite (East Sayan, Russia)
Olga N. Kiseleva 1, *, Evgeniya V. Airiyants 1 , Dmitriy K. Belyanin 1,2 and Sergey M. Zhmodik 1,2
1

2*

Sobolev Institute of Geology and Mineralogy, Siberian Branch Russian Academy of Science, pr. Academika
Koptyuga 3, Novosibirsk 630090, Russia; jenny@igm.nsc.ru (E.V.A.); bel@igm.nsc.ru (D.K.B.);
zhmodik@igm.nsc.ru (S.M.Z.)
Faculty of Geology and Geophysics, Novosibirsk State University, Novosibirsk 630090, Russia
Correspondence: kiseleva_on@igm.nsc.ru; Tel.: +7-983-126-70-72

Received: 26 December 2019; Accepted: 2 February 2020; Published: 7 February 2020

Abstract: In this paper, we present the ﬁrst detailed study on the chromitites and platinum-groupelement mineralization (PGM) of the Ulan-Sar’dag ophiolite (USO), located in the Central Asian
Fold Belt (East Sayan). Three groups of chrome spinels, diﬀering in their chemical features andphysical–chemical parameters, under equilibrium conditions of the mantle mineral association,
have been distinguished. The temperature and log oxygen fugacity values are, for the chrome spinels
I, from 820 to 920 ◦ C and from (−0.7) to (−1.5); for chrome spinels II, 891 to 1003 ◦ C and (−1.1) to (−4.4);
and for chrome spinels III, 738 to 846 ◦ C and (−1.1) to (−4.4), respectively. Chrome spinels I were formed
through the interaction of peridotites with mid-ocean ridge basalt (MORB)-type melts, and chromespinels II were formed through the interaction of peridotites with boninite melts. Chrome spinels III
were probably formed through the interaction of andesitic melts with rocks of an overlying mantle
wedge. Chromitites demonstrate the fractionated form of the distribution of the platinum-groupelements (PGE), which indicates a high degree of partial melting at 20–24% of the mantle source.
Two assemblages of PGM have been distinguished: The primary PGE assemblage of Os-Ir-Ru alloys-I,(Os,Ru)S2 , and IrAsS, and the secondary PGM assemblage of Os-Ir-Ru alloys-II, Os0 , Ru0 , RuS2 , OsS2 ,IrAsS, RhNiAs with Ni, Fe, and Cu sulﬁdes. The formation of the secondary phases of PGE occurred
upon exposure to a reduced ﬂuid, with a temperature range of 300–700 ◦ C, log sulfur fugacity of(−20), and pressure of 0.5 kbar. We have proposed a scheme for the sequence of the formation andtransformation of the PGMs at various stages of the evolution of the Ulan-Sar’dag ophiolite.
Keywords: ophiolite; chromitites; PGE mineralogy; geodynamic setting

1. Introduction
Podiform chromitites are associated with restite peridotites in ophiolite complexes [1–5].
The concentration, platinum-group element distribution, and mineral form and composition ofchrome spinels in podiform chromitites are sensitive indicators of mantle processes: the partial melting
degree, initial melt sources, and melt saturation with volatile components (S, H2 O, etc.) [6–14]. Thus, thegeochemical features of chromitites, their distribution and the mineral composition of PGE can be used
to evaluate the physical–chemical parameters, mineral equilibria parameters, initial melt composition
and geodynamic setting of ophiolite formations [14–22]. The use of an olivine–chrome spinel andolivine geothermometer, as well as experimental data on the temperature equilibria of the primary
PGMs, provides important information on the formation temperatures of the mantle olivine spinel
PGM assemblage and deformation processes [23–25]. The chemical composition of chrome spinels
and PGMs, the microstructural features of this assemblage, and accessory sulﬁde mineralization giveimportant information for the reconstruction of magmatic, hydrothermal, and metamorphic processes.
In this paper, we describe the chemistry of chrome spinels, the geochemical features of chromitites andMinerals 2020, 10, 141; doi:10.3390/min10020141

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Minerals 2020, 10, 141

PGMs from chromitites, the physical–chemical parameters of mineral equilibria, and the alteration of
PGE–chromite mineralization in the mantle peridotites of the Ulan-Sar’dag ophiolite.
2. Materials and Methods
Twenty-one samples of chromitites from the western part of Ulan-Sar’dag were studied.
The chemical composition of chrome spinels and olivines and the Al2 O3 impurities in olivine were
determined by wavelength-dispersive analysis using a Camebax-micro electron microprobe fromthe Sobolev Institute of Geology and Mineralogy, Russian Academy of Science, Novosibirsk, Russia(Analytical Center for multi-elemental and isotope research SB RAS). A JEOL JXA-8100 analysis
methodology speciﬁc to olivines to obtain their trace elements has been elaborated: The accelerating
voltage is 20 kV, the probe current is 400 nA, and the counting time per line and backgroundmeasurement are both 10 s. The number of measurements per analysis is 25 [26]. For an error of
10 relative %, the quantitation limit of Al2 O3 is within 45–100 ppm. The analytical conditions forolivines and chrome spinels are as follows: The accelerating voltage is 20 kV, the probe current is 30 nA,
the beam size is 2 μm, and the signal accumulation time is 10 s. The standards used were natural andsynthetic silicates and oxides. The detection limit for oxides was 0.01–0.03 wt. %.
The content of PGE in chromitites and dunites was determined in the rocks using the ICP-MS
microprobe and kinetic methods, along with pre-concentration (nickel matte), at the TsNIGRI analyticallaboratory (Moscow, Russia). The ICP-MS microprobe method with pre-concentration (nickel matte)
involves melting the probe (50 g) at 1100 ◦ C, during which two phases are formed: A sulﬁde (nickel
matte) and an oxide (skimming) phase. These phases are segregated according to the concentrationof all noble elements in nickel matte. The collector is an alloy of sulﬁdes and a metal alloy, in which
the bulk of the noble metals is collected [27]. The received matte is cleaned through skimming andcrushed for 10–15 s. A portion of the matte is dissolved in hydrochloric acid, while the Ni and
Cu of the matte are almost completely dissolved. Platinum metals in the nonsolute residue are
dissolved in a mixture of hydrochloric and nitric acids. The mass concentration of Pt, Pd, Rh, Ru,
and Ir in the solution is determined by ICP-MS. Osmium is separated from other noble metals by
distillation. The method is based on the ability of osmium to form volatile tetraoxides under oxidizing
conditions [28]. The determination of the osmium concentration was carried out using a kinetic method.
The method is based on the catalytic eﬀect of osmium in indicator oxidation-reduction reactions. The
reaction rate depends on the concentration of the catalyst [28–30]. The detection limits were 2 ppb for
Os, Ir, Ru, and Rh; and 5 ppb for Pt and Pd.
The studied platinum-group minerals were polished plates (21 pieces) and heavy fractions,
extracted from chromite ores. The selected PGM grains were mounted in epoxy blocks and polishedwith a diamond paste for further analysis. Microtextural observations of PGM were performed bymeans of reﬂected-light microscopy. The chemical composition of PGMs was determined using a
MIRA 3 LMU scanning electron microscope, with an attached INCA Energy 450 XMax 80 microanalysis
energy dispersive system, at the Sobolev Institute of Geology and Mineralogy, Russian Academy of
Science, Novosibirsk, Russia (Analytical Center, IGM SB RAS). Pure metals were used as the standards
to determine the chemical composition of the PGE, Ni, and Cu; arsenopyrite was used for As; and pyrite
was used for Fe. The minimum detection limits of the elements (wt. %) were found to be 0.1–0.2 for S,
Fe, Co, Ni, and Cu; 0.2–0.4 for As, Ru, Rh, Pd, Sb, and Te; and 0.4–0.7 for Os, Ir, and Pt.

32

Minerals 2020, 10, 141

3. Geological Setting
The ophiolite complexes are widely distributed in the south-eastern part of Eastern Sayan (Siberia,
Russia) (Figure 1).

Figure 1. Geological map of the southeastern part of the Eastern Sayan region [31]. NB—northernbranch, SB—southern branch.

The ophiolites are localized in the form of extended branches in the south (Il’chir) and north(Holbin-Hairhan). The features of these ophiolites are as follows: rock associations, geochemical andmineralogical characteristics, a geological structure resembling individual “massifs”, geodynamic
conditions and a formation time, which are actively studied for a considerable time. Dobretsov
and Zhmodik et al. obtained data on the heterogeneity of ophiolites of the south-eastern part ofEastern Sayan (SEPES) [32,33]. They indicate that the ophiolites of the southern branch were formed
in a mid-ocean ridges setting, and ophiolites of the northern branch were formed in an island arc
setting [32–38]. The Ulan-Sar’dag ophiolite (USO) occupies a special structural position in the Eastern
Sayan and Central Asian Fold Belt (Dunzhugur island arc). The USO is a tectonic plate, which islocated between ophiolites of the southern and northern branches of the Dunzhugur island arc, near
the contact zone of the Gargan “block” gneisses, with granites of the Sumsunur complex. The USO
is underlaid by volcanogenic and sedimentary rocks of the Ilchir suite and limestones of the Irkut
suite. Ophiolites include mantle restites (dunites, harzburgites), podiform chromitites, cumulates(metawehrlite, metapyroxenite, metagabbro), a basic dike and a volcanic complex (Ilchir suite) (Figures 2
and 3). USO has a lenticular body that is elongated in the east–west direction and is 1.5 × 5 km2 . It is
composed of dunites and harzburgites. Harzburgites predominate in the central part, dunites andserpentinites prevail in the margin, and the latter lie in contact with the volcanogenic-sedimentary
sequence of the Ilchir suite. The rocks of the cumulative series are metamorphosed under theconditions of epidote-amphibolite and amphibolite facies. The volcanogenic-sedimentary sequence33

Minerals 2020, 10, 141

is composed of volcanic, volcanogenic-sedimentary, and metasedimentary rocks (black schists andmarbles). Volcanogenic-sedimentary rocks are sulfurized, and sulﬁde mineralization is conﬁned
to the schistosity zones. High-Ti basalts (MORB tholeiites), boninites, and island-arc andesites are
distinguished in the volcanic complex. All rocks in the ophiolite nappe bottom are intensively deformed
and mark the thrust zone. They include numerous crushing zones, signs of shear displacement andgliding planes. In the contact zone of serpentinites and rocks of the Ilchir suite, the talcum powder
zones and zones of actinolite-tremolite composition are widely manifested.

Figure 2. Geological scheme of the Ulan-Sar’dag ophiolite [39] with author additions.

The types of schlieren, lenticular, and vein-like chromite bodies are localized in dunites andserpentinites (Figure 4). The length of schlieren and massive chromitite bodies ranges from a
few centimeters to several meters, and their width varies from 5 cm to 0.5 m. The predominant
structural-textural type is a massive chromitite. In some schlieren chromitites, the tectonic ﬂow
structures (Figure 4e) and the “snowball” type (Figure 4f) are observed.

34

Minerals 2020, 10, 141

Figure 3. Photographs of the ﬁeld relationship between mantle peridotites and volcanic-sedimentaryrocks (Ilchir suite), limestone thickness (Irkut suit), Gargan gneisses, and Susunur tonalities: (a) viewof the south side of the Ulan-Sar’dag ophiolite, (b) view of the north side of the Ulan-Sar’dag ophiolite.

Figure 4. Structural features of chromitite pods: (a) chromitite seams in dunites; (b) schlieren;(c) massive pods; (d) transformation of the schlieren type into the massive type due to deformation
processes; (e) folded schlieren-type chromitites; (f) structure of the “snowball”.

35

Minerals 2020, 10, 141

4. Results
4.1. Podiform Chromitites
Massive chromitites are the predominant petrographic variety of rocks in the mantle peridotitesof the USO; disseminated, schlieren, and rhythmically-banded ores are less common. Chrome spinels
make up 80–95 vol. % of massive chromitites. The intergranular space is ﬁlled with olivine andsecondary silicates: serpentine, chlorite, and, rarely, talc. The structure of massive chromitites varies
from ﬁne-grained to coarse-grained. The crystal bodies vary from subidiomorphic grains (0.1–0.5 cm)
to allotrimorphic aggregates of grains of up to 1.5 cm in size. Chrome spinels have a cataclastic texture,and they are partially fragmented. The grains are dissected by cracks ﬁlled with secondary silicates.
In the central part, chrome spinels are unchanged, and in cracks and on the grain rims, they are often
replaced by chrome magnetite. Some grains contain olivine inclusions. A wide range of minerals show
accessory mineralization in chromite ores: PGMs, sulﬁdes, sulfarsenides, and sulfosalts of base metals(Figure 5).

Figure 5. Accessory mineral association in chromitites. Hzl—heazlewoodite Ni3 S2 ; Mlr—millerite
NiS; Pn—pentlandite (Fe,Ni)9 S8 ; Cv—covellite CuS; Cct—chalcocite Cu2 S; Orc—orcelite Ni5 As2 ;Mh—maucherite Ni11As8 ; Sct—scutterudite Co4 (As4 )3 ; Apy—arsenopyrite FeAsS; Sp—sphalerit ZnS;
Gn—galena PbS; Au0 , Cu0 , Bi0 —native gold, copper and bismuth.

4.1.1. Chrome Spinels
According to the chemical composition, ore chrome spinels are divided into three groups and are
represented by chrome picotite, alumo-chromite, chromite, chrome-magnetite and magnetite (Table 1,
Figure 6). In one sample, chrome spinels of all three groups may be present. Chrome spinels ofgroups I, II, and III have the following compositions (wt. %): Al2 O3 = (17–43), (7–16), and (9–18);
Cr2 O3 = (26–54), (46–66), and (47–57); MgO = (10–20), (9–13), and (8–10); and FeO = (6–14), (1–18),
and (18–20), respectively. In chromitites of the USO, chrome picotites are found, which is typical
only of this massif, in contrast to chrome spinels from other ophiolite complexes of the southeasternpart of East Sayan [38]. Chrome spinels have a homogeneous composition. Altered chrome spinels
demonstrate increased contents of FeO, MnO, NiO, ZnO, and Fe2 O3 and decreased contents of Cr2 O3
and Al2 O3 . Chrome spinels are often replaced by chrome magnetite and magnetite along the cracks.

36

37

TiO2
Al2 O3Cr2 O3MnO
FeO
Fe2 O3
MgO
V2 O5
NiO
ZnO
Total
Al’
Cr
Fe
Mg
f
TiO2 melt
Al2 O3 melt
(Fe/Mg)mT◦ C (Ol-Sp)
f O2
T◦ C (Al in Ol)

0.1
41.0
28.5
0.5
7.9
2.3
19.8
0.8
0.3
0.0
101.1
57
40
3
71
29
0.1
18.5
0.3
893.9
−0.765

41

2

0.1
24.6
44.2
0.2
14.7
2.1
13.6
0.1
0.2
0.0
99.8
35
62
3
48
52
0.1
15.0
0.5
823.4
−1.498

74

3

0.1
24.1
43.7
0.3
14.8
2.6
13.3
0.1
0.1
0.2
99.2
34
62
3.7
47
53
0.1
14.9
0.6
820.7
−1.246

77

4

0.1
19.2
49.2
0.3
14.4
3.3
12.9
0.1
0.0
0.7
100.2
27
69
4
47
53
0.1
13.5
0.5
885.7
−1.1

101

5

0.1
7.97
63.94
19.1
0.7
10.8

234.6
26
84
0.9
50
50
0.11
9.3
0.7
991.6
−1.3

15.4
2.4
12.6

102.3
23
74
2.9
59
41
0.18
11.1
0.5
938.62
−2.48

132

7

0.10
12.19
59.57

130

6

0.1
15.3
56.4
0.3
13.3
1.2
13.7
0.1
0.0
0.0
100.4
20
80
1.7
49
49
0.1
12.3
0.4
1003.9
−2.1

120

8

0.1
14.6
54.2
0.3
14.8
2.1
12.4
0.1
0.1
0.0
98.6
20
80
3.0
54
54
0.2
12.2
0.5
948.9
−1.3

10

9

11

0.05
13.52
56.51
0.32
15.8
1.7
11.7
0.13
0.02
0.14
99.8
19
79
2.3
57
57
0.10
11.7
0.6
931.8
−1.8

48

100.7
23
74
2.9
55
55
0.13
12.7
0.5
891.6
−2.5

0.07
16.53
53.70
0.56
15.3
2.1
12.4

52

II Group

10

53

13

0.1
0.1
14.5
14.8
55.6
55.6
0.3
0.4
15.4
16.4
1.6
2.3
12.0
11.7
0.1
0.0
0.1
99.7
101.2
20
20
80
80
2.2
3.2
56
58
56
58
0.1
0.1
12.1
12.1
0.6
0.6
931.5
921.2
−1.9
−1.6
893–1332

47

12

0.1
13.9
56.5
0.3
15.6
1.7
11.9
0.1
0.0
0.2
100.2
19
80
2.3
57
57
0.1
11.8
0.6
933.2
−1.9

5

14

0.2
12.5
56.4
0.5
16.1
2.7
11.4
0.1
0.0
0.3
100.1
20
80
3.7
59
59
0.4
11.3
0.6
942.4
−1.1

112

15

9

17

0.2
0.1
12.4
14.6
57.0
57.3
0.4
0.4
16.3
17.7
2.7
1.1
11.3
10.9
0.1
0.1
0.0
0.1
0.2
0.0
100.5
102.2
20
20
80
80
3.7
1.6
59
62
59
62
0.3
0.2
11.3
12.0
0.6
0.7
913.6
861.2
−2.8
−4.4
1073–1225

88

16

0.05
14.6
54.0
0.3
19.0
1.8
9.6
0.1
0.0
0.3
99.8
20
80
2.6
66
66
0.10
12.1
0.9
812.6
−3.3

60

18

0.04
13.0
55.4
0.4
18.3
2.3
9.8
0.07
0.03
0.19
99.5
18
78
3
35
65
0.09
11.5
0.78
846.3
−3.01

71

19

21

0.07
14.0
49.0
0.4
18.6
5.6
9.6
0.11
0.06
0.10
97.6
20
71
8
34
66
0.14
12.0
0.80
821
−1.49

11

0.05
16.5
48.0
0.7
20.9
5.8
8.3
0.12
0.07
0.46
100.9
23
68
8
29
71
0.10
12.7
1.07
742.7
−1.73

22

III Group

20

0.07
19.6
48.3
0.3
19.8
2.5
9.6
0.09
0.02
0.46
100.7
28
69
4
33
67
0.14
13.6
0.96
738.1
−2.78

18

22

Notes: Al = Al/(Al + Cr + Fe3+ ); Cr = Cr/(Cr + Al + Fe3+ ); Fe = Fe3+ /(Cr + Al + Fe3+ ); Mg = Mg/(Mg + Fe2+ ); f = Fe2+ /(Mg + Fe2+ ); Sample No. 1–5—Cr-spinels I; 6–18—Cr-spinels II;
19–22—Cr-spinels III; T◦ C (Ol-Sp)—olivine-spinel geothermometer; f (O2 )—oxygen fugacity; T◦ C (Al in Ol)—olivine geothermometer.

37

0.1
40.3
29.7
0.5
7.9
2.0
19.7
0.8
0.3
0.0
101.3
56
41
2.7
72
28
0.1
18.3
0.3
919.1
−0.988N Sample

I Group

1

No.

Table 1. Representative chemical composition of Cr spinels and the calculated parental melt composition (wt. %) of Cr spinels. Thermometric and oxybarometric data
for Cr spinels.

Minerals 2020, 10, 141

Minerals 2020, 10, 141

Figure 6. Classiﬁcation diagram of chrome spinels from the chromitites of Ulan-Sar’dag, basedon the structural formula of spinels. Composition ﬁelds: I—chromite; II—alumo-chromite;
III—chrome picotite; IV—subferrichromite; V—subferrialumochromite; VI—ferrichromite;VII—subferrialumoferrichromite; VIII—chrome magnetite; IX—subalumochrome magnetite;
X—magnetite [40]. Notes: 1—chrome spinels-I; 2—chrome spinels-II; 3—chrome spinels-III; 4—chromespinels-II, with inclusions (Os,Ru)S2 ; 5—chrome spinels-II, with olivine inclusion.

4.1.2. Olivine
Olivine in chromitites occupies the intergranular space of chrome spinels; it is often replaced byserpentine or chlorite. The data, which correspond to chrysotile and forsterite, are shown in Table 2.
The MgO content is 49–54 wt. %, and the NiO variations are insigniﬁcant and amount to 0.37–0.42 wt. %,which corresponds to olivines from depleted peridotites. The fraction of the Fo component is 92–97.
Table 2. Chemical composition of the olivine from chromitites (wt. %).
N-Sample

SiO2

MgO

FeO

NiO

MnO

CaO

298-1
298-2
298-3
298-4
298-5
3-6
3-7
3-8
3-9
3-10
3-11
3-12
3-13
3-14
3-15
305-16305-172-18
973-19976-20977-21
939-22939-23970-24980-25

41.28
41.83
41.26
41.31
41.41
41.08
41.10
41.15
40.83
41.13
41.22
41.09
41.07
41.10
41.24
41.33
41.38
41.05
40.25
41.43
40.09
40.94
40.74
40.85
42.39

52.59
52.25
52.20
52.10
52.68
51.19
51.04
51.03
50.85
51.17
51.02
51.16
51.19
51.08
51.26
53.61
53.86
51.71
50.24
51.26
50.39
49.59
49.06
49.93
53.59

4.98
4.50
5.68
5.71
5.34
6.95
6.98
6.85
7.13
6.92
6.89
6.90
6.89
6.93
6.79
3.28
3.45
6.47
8.14
6.90
6.85
7.28
9.61
7.87
5.50

0.41
0.42
0.42
0.42
0.44
0.37
0.37
0.38
0.38
0.38
0.37
0.38
0.37
0.37
0.38

0.11
0.09
0.12
0.13
0.11
0.11
0.11
0.12
0.13
0.10
0.11
0.13
0.13
0.10
0.13

0.001
0.001
0.002
0.001
0.002
0.001
0.003
0.004
0.015
0.003
0.004
0.006
0.005
0.006
0.009

Al2 O30.01180.00520.019100.01190.01040.00150.01170.00050.00060.0009
0.003

0.02
0.01
0.04
0.00
0.01

Notes: (Samples 16–18)—inclusions in chromite.

38

Cr2 O3

Total

Mg#

0.002
0.002
0.004
0.002
0.002
0.002
0.001
0
0.002
0
0.001
0.001

99.38
99.08
99.68
99.68
99.98
99.73
99.60
99.54
99.34
99.69
99.62
99.66
99.65
99.58
99.80
98.83
99.38
99.23
98.64
99.61
97.35
97.85
99.81
98.67
101.51

95
95
94
94
95
93
93
93
93
93
93
93
93
93
93
97
97
93
92
93
93
92
90
92
95

0.002
0.61
0.69
0.01
0.02
0.00
0.03
0.36
0.02
0.02

Minerals 2020, 10, 141

4.2. Geochemistry of Platinum-Group Elements
The content of PGE in mantle peridotites and podiform chromitites was determined. The content
of PGE in dunites and serpentinites is 94–180 ppb; in chromitites, it ranges from 242 to 992 ppb(Table 3). The PGE content increases with an increase in the volume percentage of chromite in therock. In addition, depending on the proportion of the chromite and silicate components in the rock,the IPGE/PPGE ratio changes (Figure 7a,b) (IPGE: Os, Ir, and Ru; PPGE: Rh, Pt, and Pd). For example,
in massive ores (85–95 vol. % of chromite), IPGE > PPGE, and IPGE/PPGE = 0.86–2.15. In schlieren
lenticular chromitites and chromitites with deformational textures, the PPGE proportion increases andamounts to IPGE/PPGE = 0.03–0.67. In chromitites I, the IPGE/PPGE is higher than in chromitites II(Table 3).
Table 3. PGE abundance for the dunites and chromitites of the Ulan-Sar’dag ophiolite (ppb).
No.

1

2

3

4

5

6

7

8

9

10

11

N Sample

6

305

6 mas

7 mas

294 mas

307 mas

6 mas

3 mas

17 mas

20 mas

2 shl

Os
Ir
Ru
Rh
Pt
Pd
Total
Pt/Ir

IPGE

PPGE
IPGE/PPGEPGE
Cr# (Crt)
Al#(Crt)

6
21
7
7
35
104
180
1.67
34
146
0.23
180

8
6
14
3
15
48
94
2.5
28
66
0.42
94

45
26
68
12
39
177
367
1.5
139
228
0.61
367
68
30

51
58
53
10
49
182
403
0.84
162
241
0.67
403
67
30

58
43
116
8
31
104
360
0.72
217
143
1.52
360
65
42

117
82
221
19
54
122
615
0.66
420
195
2.15
615
57
34

49
57
121
24
49
478
778
0.86
227
551
0.41
778
70
26

37
20
59
16
64
97
293
3.20
116
177
0.66
293
69
27

81
35
44
9
46
87
302
1.31
160
142
1.13
302
84
14

46
20
46
11
41
78
242
2.05
112
130
0.86
242
83
15

7
6
20
21
35
903
992
5.83
33
959
0.03
992
75
22

Notes: 1,2—dunites; 3–11—chromitites. mas—massive pods, shl—shlieren pods; 3–6—group I chromitites;
7–10—group II chromitites; 11—group III chromitite.

Figure 7. Photographs of chromitites with diﬀerent values of IPGE/PPGE: (a) schlieren densely
disseminates in serpentinizated dunite ( PGE = 903ppb, and IPGE/PPGE = 0.03); (b) massive( PGE = 615 ppb, and IPGE/PPGE = 2.15).

4.3. Mineralogy of Platinum-Group Elements
The ﬁrst data on PGE mineralization (PGM) in chromitites of USO have been obtained.
The chemical composition of the minerals is presented in Table 4, and classiﬁcation diagrams arepresented in Figure 8a,b. Primary and secondary PGMs have been distinguished. The most common
platinum-group minerals in the chromitites of the USO are PGE sulﬁdes: laurite-erlichmanite (Ru,Os)S2 ,with diﬀerent Ru/(Ru + Os) ratios. Other PGE phases are represented by high-temperature Os-Ir-Ru
alloys-I, phases of variable-composition Os-Ir-Ru alloys-II, native Os, Ru, sulfarsenides of (Os, Ir, Ru),and zaccarinite RhNiAs. The primary and secondary PGMs will be described separately.

39

1
2
3
4
5
6
7
8
9
10
11
12
13
14
15
16
17
18
19
20
21
22
23
24
25
26
27
28
29
30
31
32
33
34
35
36

40

4.97

3.47
6.17

47.44

9.25
60.28

4.21
21.34

58.69
59.23
56.58
48.23
61.04
3.15
10.58

4.45
0
31.17
27.29
11.19
9.51
11.34
16.78
2.11
3.67

1.91
92.73
37.98
28.79
33.33
34.25
88.01
59.46
2.52
3.01

1.18
1.8

21.95
20.6
20.74
5.23
5.18
5.5
4.3
6.86
5.76
2.93
2.18
3.5

Ir

75.01
74.84
79.75
23.87
23.21
33.29
22.92
20.65
8.94
33.62
49.97

wt. %

Os

21.62
93.14
52.84
60.54
22.55
60.93
54.02
6.88
59.89
2.1
1.44
5.84
11.47
2.53
10
26.92

3.38
5.16
2.35
34.23
39.55
27.26
39.85
39.48
50.87
34.3
20.8
58.59
62.81
57.45
8.57
29.55
37.06
53.61
54.3

Ru

2.31
0.35
0.6
3.91
0.87
37.77
0.57

4.4

0.6

Rh

0.37
0.46

0.32
0.37

1.09
0.58

0.88
0.71

0,32

0.58

0.72
0.33
0.35
0.59

Fe

22.35
28.95

0.81

1.04

2.01
3.06
2.55
0.99
0.46

0.49
0.47
1.61

Ni

6.5

36.78
38.33
27.28
38.64
34.92
23.79
37.67
12.51
11.38
12.85
16.81
11.81

1.57

1.15

33.36
34.33
31.27
35.23
33,15
37.82
32
30.12
38.19
39.43
36.31
1.2

S

23.25
24.82
19.07
20.93
23.78
25.66
14.23

3.3

0.2

0.92
0.19

1.67

0.66

0

As

1.16

1

Sb

2.12
3.19
0.66

O

103.14109.09101.06100.93103.1997.28
102.8797.9
102.3
100.14103.39103.51103.07101.95102,24
101,85103.16101.1999.4
100.25102.82103.31
101.5
100.32100.5999.65
102.61
99.57
101.66
98.53
98.02
98.86
99.21
97.74
101.35

Total

0.02
0.17

0.03
0.01

0.09
0.85

0.59

0.02
0.79
0.29
0.21
0.19
0.17
0.8
0.46
0.01
0.03

0.39
0.39
0.41
0.24
0.23
0.36
0.22
0.21
0.08
0.35
0.56

apfu

Os

0.29
0.41
0.29
0.21
0.30
0.01
0.08

0.03
0.09

0.06

0.23
0.2
0.06
0.05
0.1
0.13
0.01
0.03

0.04

0.11
0.11
0.1
0.05
0.05
0.06
0.05
0.07
0.05
0.03
0.02
0.03

Ir

0.32
0.93
0.88
1
0.52
1
0.98
0.18
1.01
0.02
0.1
0.06
0.09
0.02
0.08
0.40

0.03
0.05
0.02
0.65
0.73
0.55
0.73
0.75
0.85
0,67
0.44
0.95
1.01
0,99
0.14
0.42
0.51
0.6
0.54

Ru

0.01
0.03
0.01
0.3
0.01

0.02

0.06

0.01

Rh

0.02
0.01

0.01
0.02

0.03
0.01

0,01
0,00
0.02
0.02

0.01
0.01
0.01
0.02
0.00
0.02

Fe

0,31
0.74

0.05

0.06
0.07
0.04
0.03
0.01

0.01
0.01
0.04

Ni

0.31

1.93
2
2
2
2
2
2
0.37
0.63
0.39
0.44
0.35

0.05

2
2
2
2
2
2
1.98
2.00
1.96
2
1.98
0.06

S

0.30
0.37
0.25
0.23
0.30
0.28
0.29

0.07

0.02

0.04

0.02

As

0.01

0.01

Sb

0.14
0.19

O

0.95
1
0.32
1
0.85
0.1
1

0.59
0.63
0.45
0.63
0.66
0.85
0.51
0.29
1
1
0.97

Ru/
(Ru + Os)

Notes: Primary high temperature PGM: Os-Ir-Ru alloys-I—(1–3); laurite I—(4–7); altered laurite II—(8–14); secondary PGM: Os-Ir-Ru alloys-II—(15–22); laurite III—(23–29);
irarsite—(30–34); RhNiAs—35; (Ru,Ni,Os,Ir,Rh)AsS—36.

No. anTable 4. Representative composition of platinum-group elements in the chromitite from the Ulan-Sar’dag ophiolitic massif.

Minerals 2020, 10, 141

Minerals 2020, 10, 141

Figure 8. Diagram of the compositions of PGE minerals: (a) Os-Ir-Ru alloys; (b) Laurite-erlichmanite:
Lr I—primary, inclusion in chrome spinels or isolated grains, laurite II—primary altered sulﬁdes,and laurite III—secondary sulﬁdes. The ﬁelds shown in the diagrams are drawn from [41].

4.3.1. Primary Platinum-Group Minerals
High-temperature Os-Ir-Ru alloys-I are in the form of idiomorphic inclusions in chrome spinels,
and xenomorphic grains are intergrown with laurite. Single grains contain inclusions of (Ru,Os)S2 ofless than 10 μm in size. The alloys are enriched with Os, and its content varies from 71 to 79 wt. %;the content of Ir ranges from 20 to 28 wt. %; and the content of Ru is very low and ranges from 2 to
5 wt. %. There are some impurities of Fe and Ni. In the classiﬁcation diagram, the compositions of
the primary alloys correspond to osmium with low contents of Ir and Ru (Figure 8a). Dissolution
microstructures are observed in some grains (Figure 9).
Laurite-Erlichmanite (Ru,Os)S2
Sulﬁdes can be divided into two groups, according to their composition and microstructuralfeatures (Table 4).
The ﬁrst group, laurite I, contains laurite-erlichmanite with ahomogeneous composition.
Laurite-erlichmanite occurs as an inclusion in chrome spinels (10–15 μm), and in some cases,
it can be intergrown with amphibole (magnesio-hastingsite hornblende) in chrome spinel or be foundas individual grains (Figure 9c,d). The ratio, Ru’ = Ru/(Ru + Os), is 0.61–0.78. The content of Os is
20–33 wt. %, and Ru is 27–40 wt. %. In the classiﬁcation diagram, they are in the laurite ﬁeld (Figure 8b).
The second group, laurite II, contains laurites of a heterogeneous composition, containing micro
inclusions of (Os-Ir-Ru) II alloys and laurites, which are replaced by PGE sulfarsenides (Figure 9e–g).
Laurite II (Lr II) is characterized by wide variations in the contents of Os of 1.2–49.9 wt. %, Ru of
20.8-62.2 wt. %, and the Ru/(Ru + Os) ratio of 0.44–0.99. In the classiﬁcation diagram, they are in theﬁelds of erlichmanite and laurite. In the chemical composition of laurite, a sulfur deﬁciency is often
registered (Table 4). Insigniﬁcant amounts of irarsite IrAsS are found in the chromitites of the USO,
where it replaces laurite. Irarsite forms corrosive, looped structures to replace laurite (Figure 9g,h).

41

Minerals 2020, 10, 141

Figure 9. Back Scattered Electron (BSE) images of primary PGM, showing the textural and morphologicalrelations of single and polyphase PGM from the Ulan-Sar’dag chromitites: (a) inclusion euhedral
Os-Ir-Ru alloy—I in chromite, an. 1; (b) individual single-grain Os-Ir-Ru alloys—I, with a dissolution
microstructure, an. 2; (c) inclusion in Cr-spinel of intergrowth laurite I and hornblende, an. 6;(d) individual grain of laurite I, an. 7; (e) grain of laurite II (an. 8), with inclusions of micro particle
Os-Ir-Ru alloys—II; (f) grain laurite II (an. 9), associated with an unnamed phase (Ru,Ni,Os,Ir,Rh)AsS(an. 36) (laurite II is surrounded by Os-poor laurite, which grows with millerite); (g) substitutionof laurite II by irarsite (an. 32), with remnants of laurite II (an. 10,11) and Os-Ir-Ru alloys—II;(h) laurite II (an. 12) surrounded by irarsite (an. 33). Abbreviations: Crsp—chrome spinel; Lr—laurite;
Mlr—millerite. Notes: an. No—No analysis, as shown in Table 4.

4.3.2. Secondary Platinum-Group Minerals
The secondary PGMs are presented by (Os,Ru)S2 —laurite III, native Os and Ru, Os-Ir-Ru
alloys-II of a variable composition, irarsite IrAsS, zaccarinite RhNiAs, and (Ru,Ni,Os,Ir,Rh)(As,S)
sulfarsenides of a non-stoichiometric composition. The secondary phases are localized in the chloritizedsilicate intergranular space of chrome spinels. Some grains (micro particles of osmium and otherphases) are very small (less than 2–3 μm), and their chemical composition can be determined only
semi-quantitatively. In some cases, micro particles of a variable composition (Ru,Ir,Os,Cu,Te,Ni,Ba,S,O)
can be found with secondary PGMs. Further, due to their very small size and partial coincidence withthe elemental composition of the host mineral, qualitative analysis can also be hampered.
Os-Ir-Ru Alloys-II
Micro particles of Os-Ir-Ru alloys-II are the common phases and are found in a secondary mineral
assemblage. They are localized mainly in laurite II, or they are a part of polyphase aggregates with Ni,
Cu sulﬁdes, sulfarsenides of Os, Ir, and Ru, zaccarinite RhNiAs, and laurite III (Figure 10a–d). Their
composition varies (Table 4, Figure 8a) from native osmium to native ruthenium. Native osmium iscomposed of (wt. %): Os (87–92), Ir (0–12), and Ru (3–8). Native Ru is composed of (wt. %): Ru (93),Os (2.5), and Ir (2.1). The composition of the (Os-Ir-Ru) II alloys varies, with a wide range (wt. %):Os (30–74), Ir (6–32), and Ru (8–58).

42

Minerals 2020, 10, 141

Figure 10. BSE images of secondary PGMs, showing the textural, morphological relations andassemblages of the PGM from the Ulan-Sar’dag chromitites: (a) intergrowth of laurite III (an. 26)
and PGE bearing chalcocite Cu2 S, with micro inclusions of Os-Ir-Ru alloys- II (an. 18); (b) polyphaseaggregate of agglomeration PGM particles, consisting of Os-rich laurite III (an.28), native Os0 (an. 20),and Os-poor laurite III (an. 27); (c) composite grain of irarsite (an. 34), laurite III (an. 29), and unnamed
phases (Ru, Ir, Te, Ni, S, O) within the interstitial chlorite of chromitites; (d) polyphase aggregate,consisting of Os-Ir-Ru alloys-II (an. 21), Ru0 (an. 22), zaccarinite RhNiAs (an. 35), and heazlewoodite.
Abbreviations: Chl—chlorite; Cct—chalcocite; Hzl—heazlewoodite; Irs—irarsite; Lr III—laurite III.
Notes: an. No—No analysis, as shown in Table 4.

Laurite III
The occurrence forms of secondary laurite III are very diverse. It is mainly localized in Cr-containing
chlorite, where it forms polyphase aggregates with PGE-containing chalcocite (Cu2 S) (Figure 10a),recrystallized PGM aggregates with native Os (Figure 10b), micro particles in association with
Ni3 S2 , NiS and phases of a non-stoichiometric composition (Ru,Ir,Os,Cu,Te,Ni,Ba,S,O) (Figure 10c).
The following features are characteristic of laurite III: a) A porous structure, with Cr-bearing mineralsof the chlorite group in the voids, and b) a very small grain size (micro particles). The chemicalcomposition corresponds to the end member of the laurite-erlichmanite solid solution: RuS2 -OsS2(Figure 8b). Laurite III is characterized by very low Os and Ir contents and the absence of Rh. Sometimes
it contains Ni impurities. The value of Ru varies insigniﬁcantly and amounts to 0.93–1 (Table 4);
in turn, OsS2 has low contents of Ru and Ir and is located in a recrystallized aggregate with laurite III
and native Os (Figure 10b).
The (Ir,Os,Ru)AsS in the secondary association is in the form of polyphase aggregates. It replaces
laurite. It contains Os and Ru in insigniﬁcant amounts. Secondary irarsite is presented as very small
microparticles (less than 5 μm) in chlorite in association with Ni3 S2 and laurite III.
(Ru,Ni,Os,Ir,Rh)(As,S) is found in the form of micro particles (7 μm) in a polyphase aggregate,consisting of laurite II and NiS (Figure 9f). There is a deﬁcit of S and As in this phase.
Zaccarinite RhNiAs is found in polyphase aggregates in association with Ni3 S2 , secondary(Os-Ir-Ru) II alloys and native Ru (Figure 10d).
Unknown phase No. 1 (Ir,Ni,Cu,Ru,Os,Cl) is found in the intergrowth with laurite, erlichmanite,and chalcocite. This association is localized in chromite in the cracks ﬁlled with Cr-containing chlorite.
It is worth noting the presence of Cl in this phase.
Unknown phase No. 2 (Os,Ir,Ru,As,S,O) is found in the millerite cracks in close association withchlorite and chrome-magnetite.
Unknown phase No. 3 (Ru,Ir,Te,Ni,Ba,S,O) is found in chlorite in the intergrowth with laurite andirarsite. The Ba and Te impurities are unusual for the platinum phases of chromitites.
Ni,Fe,Cu sulphides and arsenides. The sulphides of base metals form a dispersed impregnation
mainly in serpentine-chlorite aggregate. Heazlewoodite Ni3 S2 and millerite NiS are the predominantsulﬁde phases (Figure 5, Table 5). Heazlewoodite is often found in polyphase aggregates with secondary
PGE minerals. It has a homogeneous composition and is identical in individual grains and in theintergrowth with platinum metal phases.

43Minerals 2020, 10, 141

Table 5. Chemical composition of Ni, Fe, and Cu sulﬁdes and arsenides (wt. %).
No.
Mineral2a-12(Fe,Ni)9 S8

Ni
Fe
Cu
Co
Os
Ir
Ru
Rh
S
As
O
Total

38.9
25.58

6-12
Cu2 S

6-13
Ni3 S2

3-13
NiS

6-13
Ni3 S2

4-12
Ni3 S2

3-13
Ni3 As2

67.5

64.16

72.82

72.19

64.24

33.82

27.25

27.2

64.89
0.87
4.34
1.18
8.01

5.29
0.42
25.81
0.55

33.08

19.8

98.43

1.06
99.28

Incl. inolivinePGE bearing, intergrowthwith PGM

99.57

36.01
97.98

100.0799.39

100.25

PGE-free, individual grains

5. Discussion
5.1. Chromitite Formation: Composition of Parental Melts
Chromite bodies in ophiolites are formed due to the partial melting of rocks of the upper mantle.
The interaction of the melt with mantle peridotites plays an important role in the formation of podiform
chromitites. In the channel ﬁlled with molten mantle, the ascending olivine-chromite-cotectic melt
mixes with the silica-enriched melt, formed through the harzburgite-melt reaction. The formation ofsurrounding dunites along the host harzburgites is the result of a combination of olivine precipitationfrom “older” magma and the destruction of orthopyroxene in harzburgite, interacting with the magmachannel. Thin chromite schlieren and streaks can be formed by separating the chromite from thecotectic olivine-chromite melt [9,42–46].
In the Pt/Pt*-Pd/Ir diagram (Figure 11), dunites and chromitites are within the partial melting
trend. The late metamorphic and hydrothermal processes, during which Pd enrichment occurred,probably cause the high Pd/Ir ratio in some Ulan-Sar’dag chromitites. In the OSMA diagram, most of
the ore chrome spinels are in the olivine-spinel equilibrium ﬁeld, and some of the chrome spinels arebeyond this ﬁeld. This can be explained by the distortion of the magmatic system closedness, changes
in the compositions during the tectonic processes and metamorphism of chrome spinels. Chromititesof the USO were formed at the 28–35% degree of partial melting (Figure 12). The data on the dunites,
harzburgites, and chromitites of the northern and southern branches of the Ospa-Kitoy ophiolite
are presented for comparison (Figure 12). The Mg (Ol)–Cr (Sp) ratio in dunites and harzburgites
corresponds to the 30–40% degree of partial melting; in ore chrome spinels from the northern branch,it corresponds to 30–40%; in ore chrome spinels of the southern branch, it corresponds to 35%.

44

Minerals 2020, 10, 141

Figure 11. Plot of Pd/Ir versus Pt/Pt* for mantle peridotite and chromitites of the USO. The Pt anomalyis calculated as follows: Pt/[Pt]* = (Pt/8.3) × (Rh/1.6) × (Pd/4.4) [8]. US—Ulan-Sar’dag ophiolite;
ophiolite complexes from around the world: I—Wadi Al Hwanet ophiolite, Saudi Arabia [47]; II—Oman
ophiolite, Semail [48]; III—Veria ophiolite, Greece [49]; IV—Shetland Ophiolite Complex, Scotland [50];
V—Ray-Iz ophiolite, Russia [51].

Figure 12. OSMA (olivine-spinel mantle array) is a spinel peridotite restite trend [52]. The chromespinels are from: 1–4—the Ospa-Kitoy ophiolite: 1—dunites; 2—harzburgites; the chromitites are from:3—the north-branch Ospa-Kitoy, 4—south-branch Ospa-Kitoy; 5—Ulan-Sar’dag ophiolite. Mg -Mg/(Mg+ Fe) in olivine; Cr -Cr/(Cr + Al + Fe3+ ) in chrome spinel.

The composition of the chrome spinels of the USO varies signiﬁcantly, which may reﬂect theinteraction of mantle peridotites with the melts of various compositions. The joint presence of high-Cr
and high-Al chrome spinels is common in many ophiolite belts, but as a rule, chrome spinels with
diﬀerent compositions are found in diﬀerent ophiolite nappes [7,9,52–59]. Their joint occurrence in
one ophiolite nappe is less common; in this case, the ophiolite nappe, as a rule, contains peridotites
depleted to diﬀerent degrees [45,60–63]. The composition of chrome spinels of podiform chromitites
that includes, as the main components, FeO, MgO, Al2 O3 , and TiO2 is a function of the composition of
the parental melts [9,15,16,18,21,64]. We have calculated the contents of the Al2 O3 , TiO2 and FeO–MgO
45

Minerals 2020, 10, 141

ratios in the parental melts, which were in equilibrium with the podiform chromitites (Table 6),
according to Equations (1)–(3) [15]. Table 6 shows, for comparison, the values of these parameters forthe chrome spinels of ophiolites from around the world.
Al2 O3Sp (wt. %) = 0.035 × (Al2 O3 )melt 2.42 ,(1)

TiO2(melt) (wt. %) = TiO2Sp 0.82524 × e0.20203(2)

Ln(FeO/MgO)Sp = 0.47 − 1.07 × Al Sp + 0.64 × Fe3+ + Ln(FeO/MgO)melt , where Al Sp = Al/(Al+ Cr + Fe3+ ) and Fe3+ Sp = Fe3+ /(Al + Cr + Fe3+ )(3)The chemical composition of the chrome spinels I group (Crsp I) and the composition of theparental melt are similar in these parameters for medium-aluminous chrome spinels of the Ospa-Kitoy
ophiolite (Table 6). The (Al2 O3 )melt value of Crsp I is 13–18 wt. %, and the Al2 O3Sp / Al2 O3melt ratio
corresponds to the spreading trend and abyssal peridotites (Figure 13a). Despite the low TiO2 content,
Crsp I has the TiO2Sp /TiO2melt trend, which is calculated for spinels from the MORB-type peridotites(Figure 13b).
The chemical composition of the chrome spinels II group (Crsp II) and the composition of
the parental melt are similar in these parameters for the low-Al chrome spinels of the Ospa-Kitoy
ophiolite (Table 6). The (Al2 O3 )melt values for the chrome spinels of groups II and III are 10–12
and 10–13 (wt. %), and the (FeO/MgO)melt ratio is 0.4–0.85 and 0.7–1, respectively. The valuesof the Al2 O3Sp /Al2 O3melt ratios in chrome spinels II and III correspond to the trend of the chromespinels from the island-arc boninites and chromitites of the Ural-Alaska complexes (Figure 13c).
The (TiO2 )melt values for the chrome spinels I, II, and III groups overlap because of the low TiO2
content (0–0.2 wt. %). In general, the values of (Al2 O3 )melt /(FeO/MgO) are similar to the high-Cr
chrome spinels from the Troodos and Zetford Mine ophiolites (Figure 13d), which were formed in asuprasubduction setting. In the discrimination relationship diagrams, the [Al2 O3 /Fe2+ /Fe3+ ], [Mg /Cr ],and [Al2 O3 /TiO2 ] (Figure 14a–c) of the chrome spinels I group are in the ﬁeld of the MORB–typeperidotites. Chromitites were formed during the interaction of harzburgites with primitive MORB-likemelts at deep levels of the upper mantle. Through the interaction of MORB-type melts, which are in
equilibrium with abyssal dunites [65,66], precipitation of chrome spinels with Cr 0.4–0.6, as a productof the melt–peridotite reaction, is possible [43,67]. Among the volcanic rocks of the USO, metabasalts,
with enriched mid-ocean ridge basalt (E-MORB) geochemical characteristics, are available [68].
The chrome spinels II group is localized in the boninite ﬁeld. The TiO2 and Al2 O3 content
corresponds to the chrome spinels of the suprasubduction peridotites and overlaps with the chromespinels from the New Caledonia island arc (Figure 14c). The chrome spinels II are formed during thereaction of the peridotites with the island-arc boninite melt in the subduction zone. The chrome spinels
III group lies on the boundary of the boninite ﬁelds and magmatic complexes of the Ural-Alaskan type.
Three mechanisms can be suggested for the formation of the third type of chrome spinels. The ﬁrst
is the interaction with high-iron low-titanium melts [81]. The second mechanism is through plastic
deformations under mantle or crust-mantle conditions (low f O2 , Table 1), because of the reactionsof the Mg-Fe exchange with olivine. High-iron chrome spinels are found in the structural types ofchromite, including chromitites with deformation structures. When compared with chrome spinels of
the Alaskan type chromitites, the chrome spinels III group from the chromitites of the USO have low
TiO2 contents, high Cr2 O3 contents and low (mantle) f O2 values (Table 1). Variations in TiO2 contents,from 0 to 0.22 wt. %, may indicate a reaction with TiO2 -containing melts [82]. The third mechanismis through the partial melting of the ﬂuid-metasomatized mantle during the interaction of andesitic
melts with rocks of an overlying mantle wedge [71,83,84].

46

Minerals 2020, 10, 141

Table 6. Calculated parental melt composition of chrome spinels.
Cr-Spinels

Al

Cr#

Mg

Parental Melt (wt. %)References
In this article

1

I group

24–60

36–74

45–74

Al2 O3 , 13–18; TiO2 , 0–0.14
FeO/MgO, 0.2–0.7

2

II group

14–23

74–81

32–48

Al2 O3 , 10–12; TiO2 , 0–0.35
FeO/MgO, 0.4–0.85

In this article

3

III group

13–27

68–81

28–35

Al2 O3 , 10–13; TiO2 ,0.08–0.13
FeO/MgO, 0.7–1

In this article4

Ospa-Kitoy
medium Al

24—41

59–75

43–70

Al2 O3 , 12–14; TiO2 ,0.01–0.44
FeO/MgO, 0.5–1.1

[38]

5

Ospa-Kitoy low
Al

9–21

77–90

23–59

Al2 O3 , 8–11; TiO2 ,0.01–0.48
FeO/MgO, 0.5–2.4

[38]

6

MORB

35–64

29–57

57–59

Al2 O3 13–18; TiO2 , 0.3–1.7
FeO/MgO, 0.5–0.7[18,69];

7

BAB

61

34

75

Al2 O3 , 17.6; TiO2 , 0.4
FeO/MgO, 0.4

[18]

8

OIB

28

61

57

Al2 O3 , 12; TiO2 , 1.6
FeO/MgO, 0.6

[18]

9

IAB

15–29

61–68

58–69

Al2 O3 , 9–12; TiO2 , 0.4–0.7
FeO/MgO, 0.3–0.5

[18]

10

IABon, IAT

5–19

74–89

58–75

Al2 O3 , 6–10; TiO2 ,0.08–0.4
FeO/MgO, 0.2–0.4

[18]

11

LIP

13–35

52–72

35–61

Al2 O3 , 8–13; TiO2 , 0.2–0.5
FeO/MgO, 0.4–1.2

[18]

12

Abissalperidotite

45–77

20–50

64–77

Al2 O3 , 15–19; TiO2 ,0.08–0.1
FeO/MgO, 0.4–0.5

[70]

13

Chromite inophiolite mantle

40–51

44–52

63–74

Al2 O3 , 14–16; TiO2 ,0.2–0.5
FeO/MgO, 0.3–0.6

[3]

14

Chromite inophiolite mantle

21–28

67–74

56–68

Al2 O3 , 10–12; TiO2 ,0.1–0.3
FeO/MgO, 0.54–0.6[3,48]

15

Chromite in
Alaskan type

11–14

62–70

45–56

Al2 O3 , 8–9; TiO2 , 0.6–0.9
FeO/MgO, 0.4–0.5[71]

5.2. Spinel and Olivine Geothermometers and Olivine-Spinel Oxybarometers
Several processes condition the composition of chrome spinels: partial melting, cooling of mantle
peridotites and plastic deformations in the upper mantle. Many researchers have shown that, during
ultramaﬁte cooling, regardless of their formation (mantle or cumulative), (Mg ↔ Fe2+ ) exchange
reactions take place between coexisting olivines and chrome spinels, as a result of which the coeﬃcientof these elements’ distribution increases in favor of chrome spinels, and consequently, the calculated
temperatures of the olivine-spinel equilibrium decrease [85–87]. During plastic deformations, exchange
processes and rebalancing between olivine and chrome spinels also occur, and schlieren and lenticularsegregations of chromitites transform into massive chromite bodies (Figure 4c–f). Based on this, it isassumed that the obtained values for the temperature, pressure and oxygen fugacity correspond not to
the formation of ultramaﬁtes and chromitites, but to the stages of the formation and transformation
47

Minerals 2020, 10, 141of these bodies: the tectonic ﬂow of the upper mantle rocks, the eﬀect of metasomatizing ﬂuids andother processes.

Figure 13.
Plots of relations:(a) Al2 O3 _Crsp –Al2 O3_melt ; (b) TiO2_Crsp –TiO2_melt ;(c) Al2 O3_Crsp –Al2 O3_melt ; (d) calculated abundance of Al2 O3 –TiO2 in melt, as in the equilibriumwith the Cr spinels from the chromitites of the USO and the chrome spinels of chromitites from all over
the world. The ﬁelds for the chrome spinels of boninites [72–74], Troodos boninites [75,76], Thetford
boninites [77] and MORB [78–80] are shown for comparison.

We have calculated and estimated the P-T parameters with the help of an olivine-spinel (Ol-Sp)
geothermometer, provided in [88]. The oxygen fugacity was determined using an Ol-Sp oxybarometer,
provided in [89], in accordance with oxybarometers [90,91]. For the chrome spinels I group (high Al ),the calculated temperatures of the Ol-Sp equilibrium are TOl-Sp = 1020–920 ◦ C, and f O2 = (−0.7)–(−1.5)(Table 1); for the chrome spinels II group (medium-, low-Al ), the calculated temperatures are
TOl-Sp = 891–1003 ◦ C, and f O2 = (−1.1) − (−4.4); for the chrome spinels III group, these values are
TOl-Sp = 846-738 ◦ C, and f O2 =(−1.49) − (−3.01). For the chrome spinels containing inclusions oflaurite-erlichmanite, TOl-Sp = 916–938 ◦ C, and f O2 = −2.4. For the chrome spinels with the olivine
inclusion, TOl-Sp = 991.6 ◦ C, and f O2 = (−1.3). The temperature values overlap, but for the chromespinels II group, higher temperatures are noted. The values of f O2 are discrete. It is assumed thata more reducing environment and higher solid-phase reaction temperatures are required for theformation of the chrome spinels II group, in comparison with the chrome spinels groups III and I.
The content of impurities (Al2 O3 ) in olivine was used as an alternative geothermometer. For our
objects, this method was limited by the small amount of preserved olivine. We were interested in thechromitites containing PGE- and PGE-free mineralization. Using the geothermometer of De Hoogand Gall for the Al content in olivine [25], the temperature was estimated by Equation (4). For the48

Minerals 2020, 10, 141

chromitites containing PGM, the temperature, according to the olivine thermometer, was 893–1332 ◦ C,
and for the PGM-free chromitites, it was 1073–1225 ◦ C (Table 1).
Tol (C) = 1087/(7.46 − lnAlppm ) − 273(4)

An assessment of the temperature using an olivine thermometer and inclusions of primary
high-temperature (Os-Ir-Ru)-I alloys gives the values of 1200–1300 ◦ C, which is more consistent withthe expected temperatures of the chrome spinel formation from the melt.

Figure 14. Tectonic discrimination diagrams: (a) Al2 O3 –Fe2+ /Fe3+ ; (b) Mg sp –Cr sp ; (c) Al2 O3 –TiO2
from the chromitites of the USO and chrome spinels of chromitites from all over the world. The chromespinels’ composition ﬁelds for peridotites from diﬀerent geodynamic settings are drawn from [18].

5.3. Distribution of PGE in Mantle Peridotites and Chromitites
The form of PGE distribution in the mantle peridotites and chromitites of ophiolites reﬂects the
mantle conditions of chromitite formation and PGE mineralization. In the process of partially melting
the mantle source, the PGEs fractionate. The melt is enriched in Rh, Pt, and Pd (PPGE), since PPGE
is incompatible with Os, Ir, and Ru (IPGE). The mantle restite will be enriched in IPGE [17,92,93].
This process is conﬁrmed by the form of PGE distribution in the Ulan-Sar’dag peridotites (Figure 15a,b),

for which IPGE> PPGE.

49

Minerals 2020, 10, 141

Figure 15. Diagrams: (a) PGE distribution in the Ulan-Sardag peridotite literature data [94]; (b) PGE
distribution in the Ulan-Sar’dag peridotite author data; (c) PGE pattern of chromitites of ophiolitecomplexes from all over the world: The Wadi Al Hwanet ophiolite, Saudi Arabia [47]; Oman ophiolite,
Semail [48]; Veria ophiolite, Greece [49]; Shetland Ophiolite Complex, Scotland [50]; Ray-Iz ophiolite,
Russia [51]; Mongolian ophiolite [95]; 1—chromitites are enriched in Os-Ir-Ru; 2—chromitites areenriched in Pt-Pd; (d) PGE patterns in the Ulan-Sar’dag chromitites.

In the case of a low degree of partial melting, there are no obvious diﬀerences between the PPGE
and IPGE contents, and the PGE distribution is ﬂat. At a high degree of partial melting, PPGE is
depleted in restite mantle rocks, relative to IPGE, and the total PGE contents become lower. In this case,the distributions have a negative slope. The PGE distribution in the mantle peridotite of Ulan-Sar’dag
demonstrates: (1) a positive Ru and Pd picks; (2) a ﬂat type of distribution, with a negative slope
towards Pd; (3) a positive slope of Ir, Rh, and Pt and a negative slope of Ru (Figure 15b). Podiform
chromitites are characterized by a fractionated form of PGE distribution (Figure 15c), which indicates ahigh degree of partial melting (about 20%–24%) of the mantle source. Three types of PGE distribution
are observed in Ulan-Sar’dag chromitites (Figure 15d): (1) a negative slope from Os to Ir and Ru to Pt
and a positive slope from Ir to Ru and Pt to Pd; (2) a negative slope of the distribution curve fromOs to Pt and a positive slope from Pt to Pd; (3) low contents of Os, Ir and Ru and an enrichment
in Rh (Pd is uncharacteristic of chromitites). Chondrite-normalized PGE relations in chromitites of
the USO are similar to those of chromitites of ophiolites (Figure 15c,d) formed in a suprasubduction
environment from all over the world [7,11,57,63,96–98]. The pronounced positive Ru anomaly is dueto the predominant laurite phase in the chromitites of the USO. It is known that Ru has a maximum
aﬃnity with S, and the predominance of IPGE sulﬁdes, in turn, indicates a high fugacity of sulfur inthe melt, in contrast to the parental melts for chromitites of the northern and southern branches of theOspa-Kitoy ophiolite (Figure 16a–c).

50

Minerals 2020, 10, 141

Figure 16. PGE assemblage (a) Ulan-Sar’dag chromitites; (b) the northern branch of the SEPES ophiolite;(c) the southern branch of the SEPES ophiolite [36].

The experimental data show that Os, Ir, and Ru are concentrated by trapping submicroscopic
clusters of these elements in the metallic state during chromite crystallization [99–101]. Tsoupas [49]believes that an IPGE enrichment in chromitites can be associated with post-magmatic processes
during a long period of deformations, beginning from plastic deformations in the asthenosphericmantle and ending with brittle deformations in the crust. This is conﬁrmed by wide variationsin the contents of IPGE, PPGE and the IPGE/PPGE ratio from 0.03 to 2.15 (Table 3) in the studied
chromitites. Negative Pt anomalies in chromitites are closely related to the unique properties of Pt
itself. The distribution coeﬃcients of the alloy/sulﬁde liquid at 1000 ◦ C are 1–2 for Pd and more than
1000 for Pt. Thus, the distribution coeﬃcient of the alloy/sulﬁde melt for Pt is 1000 times greaterthan it is for Pd [102]. This is conﬁrmed by the distribution diagrams, where a Pt negative anomaly
is clearly visible (Figure 15c,d). An extreme Pd enrichment in one of the chromitites is most likely
associated with late magmatic processes and exposure to the reduced ﬂuid. Chrome spinels has
high FeO contents and low f O2 values. There are no Pd phases in this chromite. We believe that
PGM is concentrated in recrystallized dunite, which requires further study. Detailed studies on themetaperidotites, metagabbros, and metavolcanogenic sedimentary rocks of the USO showed thatsome rocks have signs of signiﬁcant exposure to a high-temperature ﬂuid phase, which leads to rockmetasomatism [68,103].
5.4. Sequence of the Formation and Transformation of the Platinum-Group Mineral Assemblage
Based on the chemical and textural features of PGM and associations with the magmatic andhydrothermal minerals of chromitites, several stages of PGE mineralization were distinguished(Figure 17).
•

I—Magmatic StageAt the magmatic stage, under the upper mantle conditions, euhedral–subhedral high-temperature
Os-Ir-Ru I alloys and laurite I are formed (Figures 8 and 9a–d), which are captured by chromite
grains [5,11,22,104–106]. The high osmium content in the primary (Os-Ir-Ru) I alloys is causedby an early crystallization of laurite-erlichmanite (Figure 9c). According to the experimental data,
laurite without an Os impurity crystallizes from the melt at a high temperature (T = 1200–1300 ◦ C),
P = 5–10 kbar and low log sulfur fugacity (f S2 ), from (−0.39) to 0.07 [24,106–109]. A decrease intemperature and increase in f S2 leads to a replacement of Ru by Os, and as a result, laurite rich in Os isformed. The predominance of laurite-erlichmanite over solid (Os-Ir-Ru) solutions in the chromititesof the USO is a distinctive feature, in comparison with the PGE mineralization of the northern andsouthern branches of the Ospa-Kitoy ophiolite (Figure 16a–c). In combination with the high andmedium Al chrome spinels, this indicates the formation of chromitites and PGE mineralization of
the USO because of the interaction between the initially S-saturated tholeiite magma and depleted51

Minerals 2020, 10, 141

harzburgites. Sulfarsenides and arsenides of Ru and Ir are formed from the residual ﬂuid phase at thelate magmatic stage (Figure 10g,h). With magmatic system cooling, volatile components, such as S andAs, accumulate with the formation of the residual ﬂuid phase. There is a partial replacement of laurite
by irarsite with the formation of laurite II.

Figure 17. Scheme of the alteration and transformation of platinum-group minerals.

•

II—Stage of Serpentinization and Exposure of Fluid

The PGMs in chromitites demonstrate signs of PGE remobilization (Figure 10a–d). The most intensechanges of PGMs occur at the stage of the serpentinization of ultramaﬁtes. A ﬂuid–rock interaction
occurs with the participation of reduced gases (H2 , CH4 ) and the H2 O of mantle origin. Dehydratingrocks of the subducting slab, as well as mantle-reduced ﬂuids penetrating along the fault zones intectonically weakened sectors, serve as the ﬂuid source. At this stage, the following platinum–metal
phases are formed: native Os and Ru, (Ir-Ru) alloys, phases of a variable (Os-Ir-Ru) composition,and newly formed laurite III, IrAsS, RhNiAs and RuAs (Figures 9f and 10a–d). Remobilized secondary
PGMs form polyphase aggregates in the serpentine-chlorite matrix (Figure 10c) in association with52

Minerals 2020, 10, 141

Ni sulﬁdes, sulfarsenides, and arsenides. The joint occurrence of PGMs with Ni sulﬁdes and thepresence of such elements as Ni, Fe, Te, Cu, Co, As, and Sb in the platinum phases indicate PGE
mobility in a ﬂuid-saturated medium. The processes of the redistribution and concentration of
PGE, including refractory Os, Ir, and Ru, occur at relatively low temperatures, reducing conditions
corresponding to the formation of nickel sulﬁdes, sulfarsenides, and arsenides, and low-temperature
PGE-bearing intermetallides [110]. At the initial stage, the penetration of ﬂuid through the permeablezones into laurite-erlichmanite led to the desulfurization of sulﬁdes, a deviation from stoichiometry,
the appearance of microdefects in the crystal lattice, and the formation of nanopores on the grainsurface. These processes led to the formation of microporous structures and the separation of native Os
and Ru (Figure 9e,f) [108,109,111,112]. According to the experimental data, congruent RuS2 → native
Ru decomposition occurs at T = 300 ◦ C, log f S2 = (−20), and P = 0.5 kbar [113–118]. The (Os-Ir-Ru)
phases of a variable composition are probably the products of changes in (Os,Ir,Ru)AsS, since Ir has amaximum aﬃnity with As, and irarsite therefore survives for the longest. At the same stage, secondary(newly formed) laurite III can be formed. They grow over primary laurite-erlichmanite or are conﬁned
to chloritization zones, and as a rule, they are in association with nickel and copper sulﬁdes andsulfarsenides (Figure 10a). Irarsites IrAsS can be formed during serpentinization. In this case, irarsites
are in close association with Ni3 S2 , forming joint aggregates. The microstructural features of suchaggregates indicate their simultaneous formation (Figure 10c). The physical–chemical modeling of theforms of PGE transport in the ﬂuid systems indicates the formation of carbonyl, chloride, hydrosulﬁde,and bisulﬁde complexes, in the form of which they are transported, and the formation of secondary
PGM occurs. When PGEs are transported by bisulﬁde complexes, and As and Sb appear in the system,
PGE solubility decreases, the composition of the solution changes, and the system deviates fromequilibrium. Sulfur released from bisulﬁde complexes reacts with Ni to form Ni3 S2 .
•III—Stage of Ophiolite Obduction. Regional Metamorphism.

As the ophiolite rises to the surface, the rocks undergo repeated processes of serpentinization
under the inﬂuence of metamorphogenic ﬂuids with an increased activity of O2 , As, and Sb [48].
In chromitites, chrome spinels change into chrome magnetites or magnetites. There are no clear
criteria for distinguishing remobilized PGMs under the crustal conditions (under the inﬂuence ofmetamorphogenic ﬂuids). We believe that non-stoichiometric platinum-metal phases, containing Cu,
Te, As, Sb, and O, could be formed under the crustal conditions. These elements can be transported byaqueous solutions, with a subsequent re-deposition [119,120]. These events are mainly controlled by the
Eh-pH conditions, and these minerals can be formed directly in the supergenic environment. Under theconditions of a changing temperature, varying Eh-pH in the Os-S-O-H system, low f (S2 ) and exposure
to an oxidizing high-temperature ﬂuid at a temperature of about 500 ◦ C [121], Os becomes moremobile than other PGEs, which leads to the further redistribution and redeposition of osmium. At thisstage, the following PGE-containing phases can be formed: (Ru,Ir,Te,Ni,Ba,S,O), (Os,Ir,Ru,As,S,O),and (Ir,Ni,Cu,Ru,Os,Cl). Most often, such phases are localized in the micro voids of early PGMs
(laurite, irarsite, etc.). Another process of PGM changing is the enlargement and agglomerationof PGE nanoparticles to the micro level. During progressive metamorphism (epidote-amphibolite
and amphibolite facies) and/or a thermal event (introduction of granite intrusion), the changed P-T
conditions aﬀect the stability of nanoparticles. As the temperature rises to 590–650 ◦ C, the nanoparticles
in the sulﬁde matrix become unstable [122], which leads to their coalescence (Figure 10b) andenlargement (to micron sizes). The accessory mineralization of the crust-metamorphogenic stage isrepresented by the products of the changes in Ni, Fe, and Cu sulﬁdes and sulfarsenides, with theappearance of oxygen-containing phases of a non-stoichiometric composition.
6. Conclusions(1) High- and medium-Al chrome spinels were formed through the interaction of mantle
peridotites with tholeiite melts in a spreading setting. High-Cr chrome spinels were formed

53

Minerals 2020, 10, 141

during the interaction of mantle peridotites with boninite melts in suprasubduction environments.
The predominance of PGE sulﬁdes over high-temperature Os-Ir-Ru alloys indicates their formationfrom S-saturated magma, which is typical of tholeiite melts.
(2) The formation temperatures of magmatic PGM–chromite association are estimated at
1000–1200 ◦ C. The temperatures of olivine–spinel equilibrium, reﬂecting the formation of chromitites
and tectonic deformation processes, range from 1000 to 740 ◦ C, and the log oxygen fugacity f (O2 ) is low,
ranging from (−0.76) to (−4.4), which indicates the upper mantle conditions, as well as the eﬀect of
reduced mantle ﬂuids.
(3) Platinum-group mineralization in the Ulan-Sar’dag chromitites reﬂects a long history offormation and transformation, a change in the ﬂuid conditions from magmatic to metamorphic ones.
Primary PGMs (Os-Ir-Ru alloys-I, laurite I) were formed under the condition of a high fugacity ofsulfur. The physical–chemical conditions were as follows: T < 1200–1300 ◦ C, and log f (S2 ) > (−2)/(−1).
Under the inﬂuence of reduced ﬂuids on chromitites, the desulfurization of laurite and the formationof secondary PGMs (native Os and Ru, Os-Ir-Ru alloys-II, laurite III, IrAsS, and RhNiAs) occur inassociation with serpentine, chlorite, nickel sulﬁdes and arsenides. This association can be formed at
T = 300–700 ◦ C, log f (S2 ) = (−20), and P = 0.5 kbar.
(4) At the stage of ophiolite obduction and exposure to a metamorphogenic ﬂuid, especiallyosmium, a further redistribution of PGE occurs. New phases of a non-stoichiometric composition(PGE + Cu, Te, Ba, As, Sb, O, and Cl) are formed. During progressive metamorphism, an enlargement
and agglomeration of PGE nanoparticles to the microlevel occur.
Author Contributions: Conceptualization, O.N.K. and E.V.A.; methodology, D.K.B.; software, D.K.B. and E.V.A.;
validation, S.M.Z. and E.V.A.; formal analysis, O.N.K.; investigation, O.N.K. and E.V.A.; resources, O.N.K. and
S.M.Z.; data curation, O.N.K.; writing—original draft preparation, O.N.K. and E.V.A.; writing—review and editing,
O.N.K., S.M.Z. and E.V.A.; visualization, E.V.A. and D.K.B.; supervision, S.M.Z.; project administration, O.N.K.;
funding acquisition, O.N.K. All authors have read and agreed to the published version of the manuscript.
Funding: This research was funded by the Ministry of Science and Higher Education of the Russian Federation
and Russian Fond Basic Research Grant No. 16-05-00737a and 19-05-00764a. Work is done on state assignment of
IGM SB RAS.
Acknowledgments: The work was carried out at the Analytical Center for multi-elemental and isotope research,
SB RAS. The authors are grateful of I. Ashchepkov for constructive comments and suggestions when preparing
an article.
Conﬂicts of Interest: The authors declare no conﬂicts of interest.

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minerals
ArticlePlatinum Group Elements in Arsenopyrites andPyrites of the Natalkinskoe Gold Deposit(Northeastern Russia)
Raisa G. Kravtsova, Vladimir L. Tauson, Artem S. Makshakov *, Nikolay V. Bryansky and
Nikolay V. Smagunov
A.P. Vinogradov Institute of Geochemistry, Siberian Branch of Russian Academy of Sciences, Favorskiy str 1A,
664033 Irkutsk, Russia; krg@igc.irk.ru (R.G.K.); vltauson@igc.irk.ru (V.L.T.); tridigron@yandex.ru (N.V.B.);
nicksm@igc.irk.ru (N.V.S.)* Correspondence: artem_m@mail.ru
Received: 3 March 2020; Accepted: 30 March 2020; Published: 31 March 2020

Abstract: The peculiarities of the distribution and binding forms of platinum group elements (Pt, Pd,
Ru, Rh, Os and Ir) in the arsenopyrites and pyrites of the Natalkinskoe gold ore deposit (Northeastern
Russia) were examined using atomic absorption spectrometry with analytical data selections for single
crystals (AAS-ADSSC), a “phase” chemical analysis (PCA) based on AAS of diﬀerent size-fractions ofminerals, scanning electron microscopy with energy dispersive X-ray spectrometry (SEM-EDX) andlaser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS). The arsenopyrites andpyrites of the Natalkinskoe gold deposit were found to concentrate not only Au but also platinumgroup elements (PGEs) such as Pt, Pd, Ru and Rh. The PCA showed that the highest contents (in ppm)
were found in the monofractions of arsenopyrite—Pt up to 128, Pd up to 20, Ru up to 86 and Rh up to
21—and comparably lower in monofractions of pyrite—Pt to 29, Pd to 15, Ru to 58 and Rh to 5.9.
The AAS-ADSSC method revealed two forms of uniformly distributed Pt, Pd and Ru corresponding
to the chemically bound element in the structure of the mineral and in the superﬁcial non-autonomous
phase (NAP). The superﬁcially bound form dominates over the structural form and presumably
exists in a very thin surface layer of the crystal (~100–500 nm). The maximum contents of these PGE,
chemically bound in the structure of arsenopyrite, reached values of (in ppm) 48, 5.9 and 48; and in
pyrite structure, 68, 5.2 and 34 for Pt, Pd and Ru respectively. The contents of Pt, Pd and Ru related to
NAP on the surface of the crystal were signiﬁcantly higher and amounted (in ppm) for arsenopyrite
to 714, 114 and 1083; and for pyrite 890, 62 and 690 for Pt, Pd and Ru, respectively. Preliminary results
for the Rh form in arsenopyrite crystals suggest that the surface-related form (154–678 ppm) is
more abundant than the structural form (17–45 ppm). Data from studying the surfaces of sulphideminerals by SEM-EDX and LA-ICP-MS conﬁrmed the presence of Pt, Pd, Ru and Rh on the surface
of arsenopyrite and pyrite crystals. These methods generated primary data on the content of Os
and Ir in arsenopyrite and pyrite in the surface layer. The maximum content of Os and Ir foundin arsenopyrites was up to 0.7 wt%. PGE-enriched ﬂuids (up to ~3 ppm Pt) may exist in the goldore deposit. It is assumed that there is a common mechanism of impurities uptake associated withthe active role of the crystal surface and surface defects for gold-bearing arsenopyrites and pyrites.
The surface enrichment is due to peculiarities in the crystal growth mechanism through the mediumof NAP and the dualism of the element distribution coeﬃcient in the system of mineral–hydrothermal
solution, which is higher for NAP, compared to the volume of the crystal. Although mineral formsof Pt, Pd, Ru, Rh, Os and Ir have not been found at the Natalkinskoe gold deposit, their existencein the form of nano-scale particles is not excluded. This follows from the evolutionary model of
surﬁcial NAPs, assuming their partial transformation and aggregation with the formation of nanoand micro-sized autonomous phases of trace elements. The presence of PGE in the ores and thepossibility of their extraction signiﬁcantly increase the quality and value of the extracted raw gold
materials at the Natalkinskoe deposit, and adds to the list of known platiniferous ore formations.
Minerals 2020, 10, 318; doi:10.3390/min10040318

61

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Minerals 2020, 10, 318

Keywords: Natalkinskoe gold deposit; arsenopyrite; pyrite; platinum group elements; distribution;binding forms

1. Introduction
Studying the platinum content of gold ore deposits and platinum group elements (PGEs) ingeneral in the rocks of carbon-containing and black shale formations is currently a promising areaof research, both in Russia and abroad. Many researchers see these deposits as a promising newmineral resource, although views on the presence and potential of extracting PGE from the ores of these
deposits are somewhat contradictory [1–13]. The Natalkinskoe deposit (Northeastern Russia, Magadan
region) is one of the largest gold deposits in Russia, and so is of particular interest. In addition togold, there are quite high contents of Pt and Pd in the ores of the deposit. There are few data on other
PGEs (Ru, Rh, Ir and Os). PGE contents and distribution studies are necessary to complete the mostimportant tasks in the geochemistry of endogenous ore formation, such as the study of the material
composition and the identiﬁcation of genesis features of deposits, including giant and unique deposits.
These studies are of great practical importance. The presence of Pt in ores and the possibility of its
extraction would signiﬁcantly complement the list of already known platiniferous ore formations andsigniﬁcantly increase the quality and value of the gold ore raw materials mined.
Work to detect platinum content at the Natalkinskoe deposit began in 1990 by a research groupheaded by V.I. Goncharov from the North–East Interdisciplinary Scientiﬁc Research Institute of the
Far-East Branch of the Russian Academy of Sciences (Magadan, Russia) [1–3]. Much geochemical
and mineralogical work has been performed over the years, but in terms of contents of PGE “ . . . theobtained analytical results remain ambiguous, and often contradictory” [3] (p. 132). The reliabilityof previously obtained high Pt content in ores (tens of ppm) and Pd (several ppm) [3,14] is stillquestionable. It was concluded that “Minerals-concentrators of platinum group elements are not foundin the deposit” [3] (p. 136) nor in native mineral forms of PGE.
Subsequently, and until now, questions relating to this issue have been raised repeatedly, but there
has been no clear response. Plyusnina and co-authors [14] (p. 840) concluded that “ . . . the industrial
content of precious metals (Au and Pt) at the Natalkinskoe deposit is caused by metasomatic alterationand sulphidization of carbonaceous rocks”. Unlike Au, the involvement of the carbonaceous substancein the accumulation of Pt has not been established by the authors of this work: the connection between
this metal and the organic component of rocks and ores has not been found. It was assumed thatsulphide minerals concentrate not only Au [3,15], but also Pt [14]. Later arsenopyrite was foundto act as an Au, Pt and Pd concentrator at the Natalkinskoe deposit [11]. According to preliminary
data obtained at that time, Pt contents in monofractions of this sulphide ranged from 23 to 62 ppm,
and Pd from 2.3 to 9.5 ppm. Two main non-mineral forms of Pt and Pd, structural and surface-related,were identiﬁed. The nature of high contents of non-mineral Pt and Pd forms in arsenopyrite was foundto be mainly superﬁcial. Distinct mineral forms of Pt and Pd were not observed.
Considering all the above, the study of the binding forms and concentration levels of PGE in
the most common sulphide minerals at the Natalkinskoe deposit, arsenopyrite and pyrite, seems to
be an interesting and important task. This paper presents new data on the content of Pt and Pd inarsenopyrite and pyrite. The concentration levels of Ru, Rh, Ir and Os in these minerals are estimatedfor the ﬁrst time. It is shown that the ambiguity of the results of previous studies may have beencaused by the presence of hidden (so-called “invisible”) forms of PGEs.
2. Research Objects
The Natalkinskoe gold ore deposit, where this research was conducted, is located in the territoryof Northeastern Russia (see Figure 1, inset). It is part of the Omchak ore-placer cluster and one of
the largest gold reserves in Russia. Structurally, the deposit is conﬁned to the boundary part of the62Minerals 2020, 10, 318

estimated granite pluton in the Tenka deep fault zone and is associated with the collision development
stage of the Yana–Kolyma fold system. The primary objects of the study are arsenopyrites and pyrites
selected from the most industrially signiﬁcant vein, streaky-vein and veinlet-disseminated ores of
the Natalkinskoe deposit, which belong to the productive gold-bearing quartz–pyrite–arsenopyrite(with native gold) mineral association.
The deposit falls under the category of orogenic gold–quartz ore formations, and it ischaracterized by a complex, long-term nature of development and, according to most researchers,by metamorphogenic-hydrothermal genesis. The ore hosted rocks are identiﬁed as late Palaeozoic(upper Permian), and ore mineralization is presumably Mesozoic (from late Jurassic to late
Cretaceous) [3]. The ores of the Natalkinskoe gold deposit, unique in scale, despite the diversity,
form a deposit with uniform internal structure consisting of zones of quartz, carbonate–quartz,
sulphide–quartz veins and veinlets, surrounded by a wide halo of sulphidized rocks. They are
generally characterized by the same elemental and mineral composition, diﬀering only in the degreeof their manifestation: the quantitative ratios. The host volcanogenic sedimentary strata include
diamictites, argillites and siltstone shales with layers of sandstones and gravelites with high carbon
content. The ores of the deposit, as already mentioned, demonstrate relatively high contents of Pt
and Pd. Commercial ores emerged in the interaction of host rocks with low- and moderate-salinity
water–bicarbonate ﬂuids in the salinity range of 3–12 wt% eq. NaCl, at temperatures of 360–280 ◦ C andpressures of approximately 2.4–1.1 kbar [16]. More details on the geology, mineralogy, geochemistry
and conditions of formation of the Natalkinskoe giant gold deposit are available [1–4,11,14–28].
To study the distribution features and forms of occurrence of PGEs in arsenopyrites and pyrites,
35 large-volume ore samples were selected, each weighing up to 10 kg. The samples were taken fromquarries and boreholes of diﬀerent horizons of the deposit from those most rich in the number of sulphidemineral veins, veinlets and metasomatites classed as gold vein, streaky-vein and veinlet-disseminatedores (the areas named “Geological”, “North–West”, “Central” and “South–East”). The schematic
geological map of the Natalkinskoe gold deposit with areas from which samples were taken is shownin Figure 1, ore types in Figure 2.
Non-metallic minerals such as quartz, carbonates, feldspar, sericite, chlorite and carbonaceous
matter (micrographite) were found in the specimens characterizing these samples. Arsenopyrite andpyrite amounted to 95–99% of the ore minerals, with the former predominating. Ore minerals,
along with arsenopyrite and pyrite, the amount of which in some areas reached ~7%, included galena,
sphalerite, chalcopyrite, rutile and native gold (up to 1%). Native gold with a ﬁneness from 750%,to 900%, less frequently electrum, mainly of large grain size, was located mainly in the vein quartz
and intergrowths with sulphide minerals. Small quantities of pyrrhotite, sulphoarsenides of nickel andcobalt and ilmenite were identiﬁed. More than 70 minerals were found at the deposit in total [3,19,20].

63

Minerals 2020, 10, 318

Figure 1. The schematic geological map of the Natalkinskoe gold deposit with areas from which ore
samples were taken. Constructed by the authors of the present paper using the data [3,19,21,24,26].
Legend: 1—Quaternary sediments; 2 and 3—upper Permian ore hosted volcanogenic sedimentaryrocks: 2—diamictites, 3—argillites and siltstone shales with layers of sandstones and gravelites withhigh carbon content; 4—ore deposit; 5—faults (a—reliable, b—estimated); 6—sampling points andtheir numbers. The inset map shows the location of the Natalkinskoe deposits in the territory ofNortheastern Russia.

64Minerals 2020, 10, 318

Figure 2.
Types of ores in the Natalkinskoe deposit (a–c—vein, d–f—streaky-vein,
g–i—veinlet-disseminated): (a) a massive quartz vein with a small number of fragments containing
siltstones, nests and inclusions of arsenopyrite (”North–West” area, borehole DH-20/11n, interval
566–569 m); (b) quartz vein of laminated texture (bands are composed of completely siliciﬁcatedsiltstones) with the impregnations of small grains of arsenopyrite (”Central” area, quarry, elevation
810 m); (c) massive quartz–sulphide vein of breccia structure with fragments of host diamictites andenclaves of arsenopyrite and pyrite crystals (“North–West” area, quarry, elevation 900 m); (d) siltstones
dissected by a net of quartz and carbonate–quartz veinlets with nests and individual crystals ofarsenopyrite and pyrite, on the border with a massive quartz vein (”Central” area, mine, elevation
600 m); (e) siliciﬁcated diamictites dissected by quartz and carbonate–quartz veins with enclaves ofarsenopyrite and pyrite crystals on the border with a massive quartz vein (“South–East” area, borehole
DH50/12n, interval 452.0–454.9 m); (f) massive quartz vein on the border with diamictites, dissected by
the carbonate–quartz veinlets, with fragments of sandy argillites of gravel size with enclaves of crystals
and clusters of arsenopyrite and pyrite (“South–East” area, borehole DH50/12n, interval 460.8–463.8 m);(g) diamictites with fragments of argillites of gravel size dissected by a dense net of quartz andcarbonate–quartz veins, with enclaves of arsenopyrite and pyrite crystals (”Central” area, borehole
DH30/29, interval 273.5–276.5 m); (h) diamictites with enclaves of crystals, clusters and crosscuts ofarsenopyrite and pyrite, from the zone of intense quartz and quartz–carbonate streaking (“South–East”
area, borehole DH50/12n, interval 428.1–431.0 m); (i) diamictites dissected by transverse quartz andquartz–carbonate veins with fragments of argillites of gravel size and enclaves of arsenopyrite andpyrite crystals and clusters (“South–East” area, borehole DH50/12n, interval 528.4–531.4 m).

65

Minerals 2020, 10, 318

3. Research Methods
All ore samples were analyzed for Au and PGEs using atomic absorption spectrometry (AAS)
and inductively coupled plasma mass spectrometry (ICP-MS) [29–31]. Twenty-two monofractions ofarsenopyrite and 14 monofractions of pyrite were selected from the 7 ore samples richest in amount ofsulphides, Au and PGE. They were all subjected to AAS analysis for Au, as were the ore samples [29,30].
All sulphide minerals were found to be gold-bearing. The level of Au content in arsenopyrite was
1.4–1383 ppm, in pyrite 0.8–158 ppm.
Further study of these gold-bearing arsenopyrites and pyrites for PGE was carried out using“phase” chemical analysis (PCA) based on AAS, coupled with analytical data selections for single crystals(ADSSC) [32,33], scanning electron microscopy with energy dispersive X-ray spectrometry (SEM-EDX)
and laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS). These methods best
characterize the peculiarities of minerals such as internal parts and surface layers, microinclusions,the so-called “invisible” impurity elements, their forms, distribution and concentration levels.
This set of methods was applied because PGEs in many deposits are represented by at least two
forms—a solid solution in a carrier mineral and the submicroscopic inclusions of a PGE-rich material
in an invisible form, undetectable by optical and scanning electron microscopy [34]. The results of theexperiment and theoretical analysis suggest that not only Au, but also other noble metals (PGE and Ag)
can form chemically bound forms of impurities in pyrite and other sulphides and chalcogenides [35].
The presence of these forms (at least one of them—a solid solution of PGE in sulphide or sulphoarsenide
matrix) is identiﬁed even in low-temperature (~100 ◦ C) platinum-bearing associations [36].
Current methods of “phase” chemical analysis for the determination of PGE in monofractions ofarsenopyrite and pyrite, as well as the analysis of individual crystals of these minerals by AAS-ADSSC,
allow a reliable determination of the contents of only three PGEs: Pt, Pd and Ru. The Rh analysis
is under development, and the data are preliminary. All PGE—Pt, Pd, Ru, Rh, Ir and Os—can beanalyzed using SEM-EDX and LA-ICP-MC methods. Restrictions are related only to their detection
limits. Requirements for the selection of the test material are discussed below.
Monomineral samples of diﬀerent sizes (0.07–0.14, 0.14–0.2, 0.2–0.25, 0.25–0.5, 0.5–1.0 and
1.0–2.0 mm) are taken for “phase” chemical analysis: individual and cluster crystals of arsenopyrite
represented by pseudo-orthorhombic and monoclinic prisms; individual and cluster crystals of pyriteof cubic, pentagon-dodecahedral and cuboctahedral habits (Figure 3).

Figure 3. Monomineral fractions: (a) arsenopyrite, (b) pyrite.

In order to study PGEs in arsenopyrites and pyrites using the AAS-ADSSC method, euhedral
individuals with well-deﬁned morphology were selected from monomineral fractions of diﬀerent sizes,
from 0.25 to 2 mm. The crystals of arsenopyrite were predominantly shaped as pseudo-orthorhombic

66

Minerals 2020, 10, 318

prisms, close to a slightly distorted parallelepiped, and less frequently as monoclinic prisms.
Selected pyrite crystals generally had the shape of a cube or parallelepiped, although the samples
were often complicated by pentagonal dodecahedron and octahedron facets (Figure 4). Shape is thecharacteristic required for calculation of the speciﬁc surface area of average crystal in size sample
using the form coeﬃcient for the true polyhedron. In our case, the coeﬃcient was six for cubes andparallelepipeds. A total of 415 arsenopyrite crystals and 234 pyrite crystals were selected and studied.

Figure 4. Crystals: (a) arsenopyrite, (b) pyrite.

Twenty arsenopyrite crystals and 14 pyrite crystals were selected and examined by SEM-EDX and
LA-ICP-MS in order to study the PGE distribution in the surface layer. Euhedral pseudo-orthorhombic
crystals of arsenopyrite and cubic crystals of pyrite 0.5–1 mm in size were used to study the surfaceof these minerals by SEM-EDX. Larger (1–2 mm) crystals of arsenopyrite and pyrite with at least
one clearly manifested face belonging to certain simple form were selected for the LA-ICP-MS study(see Figure 4).
3.1. “Phase” Chemical Analysis
PCA was carried out with 10–20 mg samples with the view to determine the PGE in monofractionsof arsenopyrite and pyrite. Monofractions were ground to powder and decomposed with Aqua
Regia during heating. After decomposition, the material was treated with concentrated HCl andevaporation to dry state to remove nitric acid residues and convert the salts to chloride form.
After cooling, the samples were diluted with 2M HCl to the condition required for the AAS methodwith electrothermal atomization on the M-503 device (Perkin-Elmer, Waltham, MA, USA) usinggraphite atomizer HGA-72 (Perkin-Elmer, Waltham, MA, USA) [29,30]. Au was determined in all thesamples, along with Pt, Pd and Ru [29]. The accuracy of determination was ±10–12 rel. %.
3.2. AAS-ADSSC Method
The study of the so-called “invisible” uniformly distributed impurity component of PGE inarsenopyrites and pyrites was carried out using the AAS-ADSSC method. The method was developedby V.L. Tauson and co-authors for the study of structural and surface-bound forms of gold in ore
minerals [32,33,37,38]. The method was further reﬁned and successfully used to study structural andsurface-bound forms of Pt, Pd, Ru and Rh [11,38,39]. The technique is designed to determine the content
of the structural impurity of the element with an uncertainty of ±30 rel. % [40,41]. The determination ofPt, Pd, Ru and Rh in solutions obtained through the decomposition of individual crystals of arsenopyrite
and pyrite was carried out following their preliminary extraction, concentration and separation fromthe matrix. Tristyrylphosphine, (C6 Н5 CН–CH)3 P, was used as an extracting agent. Extraction was
carried out using hydrochloric acid solutions (0.5 M HCl). The extracting agent concentration was67Minerals 2020, 10, 318

0.05 M (in toluene), and the phase contact time was 30 min. The ratio of aqueous to organic phases was
2:1. The extraction was carried out in static mode at room temperature and without labilising additives.
The organic phase was used to measure element concentrations. Measurements were performed by
AAS with electrochemical atomization on an М-503 (Perkin-Elmer, Waltham, MA, USA) device withgraphite atomizer HGA-72 (Perkin-Elmer, Waltham, MA, USA) [29,30].
The data obtained were statistically processed according to the regulations of the distribution of
diﬀerent binding forms of the element [33,38]. We divided the dataset (>30–40 crystals with evenly
distributed element) into the intervals of crystal masses (sizes), chosen to be as narrow as possible,
although statistically representative, and determined the average crystal mass in every size fraction(m), average size (r), speciﬁc surface area (Ssp ) and element content (C). The more size fractions andnumber of crystals in the ﬁnal samples, the more reliable the results obtained. When constructing thedependences C = f (Ssp ), we usually obtained a number of points best approximated with an exponent.
The extrapolation of these curves to a zero-speciﬁc surface, that is, to a virtual inﬁnite crystal, gave
the bulk element content. In our model, this was equal to the structurally bound element content(Cstr ) because all other possible bulk forms were eliminated at the stage of initial dataset processing.
The superﬁcially bound element content (Csur ) characterizes an average crystal among all size samples,that is, the surface-related excess content of the element, and can be calculated with the equation given
in [38]. It is important to note that in such an approach, the amount of the material is normalized to thewhole crystal, and this allows a comparison of the contribution of surface and bulk related forms of theelement to its total content.
3.3. SEM-EDX Method
The study of the “invisible” impurity content of the element on the surface of arsenopyrite
and pyrite crystals by SEM-EDX was carried out on a Quanta 200 scanning electron microscope(FEI Company, Hillsboro, OR, USA) with energy-dispersive accessory EDAX (FEI Company, Hillsboro,
OR, USA) with nitrogen-free cooling for X-ray spectral microanalysis. The maximum spatial resolutionof the device with a tungsten cathode and a standard detector of secondary and backscattered electrons
was 3.5 nm. In elemental analysis, the resolution maximally amounted to 2 μm. The EDAX accessory
with EdaxGenesis software enables not only semiquantitative but also quantitative analysis of elements
in a wide range (from Be to U) with a resolution of 127 eV and a minimum detection limit of 0.5 wt%
for the elements considered. Microphotography and crystal surface analyses were performed underhigh-vacuum conditions, without sputtering and at an accelerating voltage of 30 kV. Elemental analysis
was performed on a ﬂat area with a side of at least 2 μm. The contents were calculated using thethree-correction ZAF method (atomic number, absorption and ﬂuorescence). The ZAF method is basedon measuring the X-ray intensity of the i-th element of the sample (Ji ) and the standard of the known
composition containing this element (J(i) ) under the same conditions, minus the intensity of the X-ray
background radiation, followed by normalization by 100%. To study the chemical composition of thesurface of arsenopyrite and pyrite crystals using the SEM-EDX method, the most “clean” areas, with
virtually no inclusions or signiﬁcant defects visible at maximum magniﬁcation were selected. It seems
meaningless to use polished samples while examining natural surfaces. However, it is possible tominimize ﬂuctuations of the probe current by selection of smooth-faced sections and their positioning
in the SEM device. Sections with as little roughness as possible (2–3 nm, well comparable withordinary polished surfaces) were selected using atomic force microscopy on a SMM-2000 scanningprobe microscope (Proton-MIET, Moscow, Russia).
3.4. LA-ICP-MS Method
The study of the surface layer of arsenopyrite and pyrite crystals was continued using the
LA-ICP-MS method. A quadrupole inductively coupled plasma mass spectrometer NexION 300D
(Perkin-Elmer, Waltham, MA, USA) was used in combination with a laser ablation system NWR-213(New Wave Research, Fremont, CA, USA) with a laser wavelength of 213 nm. The working and carrier

68Minerals 2020, 10, 318

gases were Ar and He, at 99.999% purity. The laser burned a continuous groove (track) no longer
than 1.4 mm on the surface of the crystal. The diameter of the laser beam was 100 μm, the frequency
was 10 Hz and the energy on the surface of the sample was 0.4 J/cm2 . The laser movement speed
was 70 μm/s, and the background measuring time was 20 s. The calibration graphs were constructedin compliance with the NISTSRM610, NISTSRM612, NISTSRM614, BHVO-2G, TB-1G and NKT-1G
international standards. It was possible to verify the results for several elements (Au, Pd and Pt) using
the in-house sulphide reference sample, which included highly homogeneous ferrous greenockite(α-CdS) crystals synthesized hydrothermally at 500 ◦ C and 1 kbar pressure [42]. The instrument error
determined for PGE did not exceed 10%. The contents of the elements in arsenopyrite and pyrite
were calculated according to the depth of the track. The ablation depth was controlled by a Quanta
200 scanning electron microscope (FEI Company, Hillsboro, OR, USA). An analysis depth from 0.5 to
1 μm per laser pass was obtained for all samples. For one of the arsenopyrite crystals (sample M-129/10)
selected for a detailed study, the laser parameters were adjusted in such a way as to achieve the burned
layer of ~100 nm in depth in one pass. Depth was controlled using atomic force microscopy with a
SMM-2000 scanning probe microscope (Proton-MIET, Moscow, Russia). In this case, the diameterof the laser beam was 5 μm, the frequency 20 Hz, and the energy on the surface of the sample was
10 J/cm2 with similar parameters. The instrument error for Pt and Pd did not exceed 30%.
4. Results and Discussion
This section presents data for the study of distribution features and forms of occurrence of PGE(Pt, Pd, Ru, Rh, Ir and Os) in arsenopyrite and pyrite from the Natalkinskoe gold deposit. The data
were obtained through local (SEM-EDX and LA-ICP-MS) and, so to speak, “semi-local” analyses (PCAand ADSSC). That is, the methods are not exactly local but were by no means bulk analysis, with aresolution at the level of diﬀerent size fractions (PCA) or the size of individual crystals (ADSSC).
The combination of local analysis and statistics for the compositions of individual monocrystals of
diﬀerent sizes is a unique feature of our approach.
4.1. Pt, Pd, Ru and Rh Content in Monomineral Fractions of Arsenopyrite and Pyrite According to PCA
All studied arsenopyrites and pyrites from the deposit were found to be concentrators not only ofAu, but also of such elements as Pt, Pd and Ru. Data on their content are given in Table 1. The highest
contents (in ppm) were observed in arsenopyrite: Pt up to 128, Pd up to 20 and Ru up to 86, with lower
values identiﬁed in pyrite: Pt up to 29, Pd up to 15 and Ru up to 58. The ﬁrst, although preliminary,
results on the content in the monofractions of the studied sulphides of another platinum group element,
Rh, have been obtained (Table 2). Maximum contents of this element, up to 21 ppm, as well as for Pt,
Pd and Ru are found in arsenopyrite, and lower levels up to 5.9 ppm were found in pyrite. Not only
gold-bearing arsenopyrites, but also pyrites were shown to act as concentrators of these elements.
It can be argued that the most platinum-rich ores have a sulphide–quartz composition. The data on
PGE content in bulk ore samples are in favor of this argument. Relatively high contents (in ppm) wereobserved for Pt (up to 0.260) and Pd (up to 0.029), and elevated contents for Os (up to 0.003), Ru (up to
0.004), Rh (up to 0.0012) and Ir (up to 0.0006).

69

70

Nat-1019

22

21

UV-3/13

M-161/10

15
16
17
18

20

M-131/10

M-129/10

TPM-1/1

G-9/13Sample
No.*

12
13
14

10
11

9

7
8

6

1
2
3
4
5

FractionNo.

0.25–0.5

Streaky-vein ore type. Veins and veinlets of quartz and carbonate–quartz indiamictites, with inclusions of crystals and nodules of arsenopyrite, lessfrequently—pyrite. The crystals of arsenopyrite are shaped as
pseudo-orthorhombic and monoclinic prisms.

0.2–0.25 **Veinlet-disseminated ore type. Diamictites interspersed with pyrite, rarely
arsenopyrite, dissected by a dense net of quartz and carbonate veins fromhair-thick to 5 mm. The crystals of arsenopyrite are shaped as
pseudo-orthorhombic and monoclinic prisms.

Veinlet-disseminated ore type. Diamictites interspersed with arsenopyrite andpyrite, dissected by a dense net of quartz and carbonate veins. Arsenopyrite
crystals shaped as pseudo-orthorhombic and monoclinic prisms predominate.

27
47
62
44

1–2
0.5–1
0.25–0.5
0.14–0.25

Vein ore type. Massive sulphide–quartz vein with a small number of fragmentsof diamictites, with nodes and inclusions of arsenopyrite, rarely pyrite.
Arsenopyrite crystals shaped as monoclinic prisms dominate.

0.2–0.25

0.25–0.5 **

0.5–1

24
26
29

0.5–1
0.25–0.5
0.14–0.25

23

3.5
0.7
0.4
0.7

1.9

20
15
5.9
8.0
10

3.4

3.7
6.3
9.5
6.2

2.8
2.3
3.6

3.9
4.1

3.5

11
6.7

4.0

12
20
5.4
14
17

Pd
(ppm)

20

25
27

Streaky-vein ore type. Oﬀset subparallel quartz and carbonate–quartz veins andveinlets in diamictite. Sulphides are mainly represented by arsenopyrite. The
crystals are shaped as pseudo-orthorhombic and monoclinic prisms.

0.5–1
0.25–0.5
0.2–0.25

73
120

40

60
70
128
64
63

Pt
(ppm)

Vein ore type. Sulphidized quartz vein on contact with diamictites dissected bya dense net of quartz and carbonate–quartz veins. Sulphides are mainly
represented by arsenopyrite dominated by crystals in the form ofpseudo-orthorhombic and monoclinic prisms.

0.14–0.25
0.07–0.14

0.5–1
0.25–0.5
0.2–0.25
0.14–0.2
0.07–0.14

Fraction(mm)

Veinlet-disseminated ore type. Diamictites dissected by a dense net of quartz
and carbonate–quartz veinlets with inclusions of arsenopyrite and pyrite
crystals. Arsenopyrite crystals shaped as pseudo-orthorhombic and monoclinic
prisms predominate.

Arsenopyrite

Characteristics of Samples for Selection of Monofractions

3.7
5.4
4.0
4.2

17

20

19
33
31
28

16
24
17

21
24

18

54
67

31

44
55
25
53
86

Ru
(ppm)

Table 1. Platinum, palladium and ruthenium content in diﬀerent-sized monofractions of arsenopyrite and pyrite selected from vein, streaky-vein andveinlet-disseminated ores of the Natalkinskoe deposit.

Minerals 2020, 10, 318

71

36

0.2–0.25

0.25–0.5 **

12
15
14
4.7
1.0
1.0
1.0
2.1
5.8
2.8
3.6
4.9
1.0
0.5
1.0
1.0
1.7

25
10
15
15
12
10
10
13
18
15
8.0
11
17
29

0.5
0.3
0.4

Pd
(ppm)

14
25
27

1.0
3.3
1.1

Pt
(ppm)

16
2.8
28
16
14

21
28
57
39
58
48
22
11
17

24
33
41

4.4
2.9
4.2

Ru
(ppm)

Note: The data for the “phase” chemical analysis are presented here and in Table 2. Here: elev.—elevation relative to sea level. * Ore sampling location: G-9/13—“Geological”
area, borehole DH329n, interval 151.6–154.6 m (elev. 750 m); TPM-1/1—“North–West” area, quarry (elev. 860 m); M-129/10—“North–West” area, ore zone 33, surface (elev. 920m); M-131/10—“North–West” area, ore zone 33, surface (elev. 900 m); M-161/10—“Central” area, mine (elev. 600 m, crosscut 11); Nat-10—“Central” area, quarry (elev. 790 m);
UV-3/13—“South–East” area, borehole DH70/5n, interval 160.1–163.1 m (elev. 590 m). ** Measurements of Pt, Pd and Ru were carried out on two parallel samples.

UV-3/13

0.5–1 **

34

35

0.14–0.2

Veinlet-disseminated ore type. Diamictites interspersed with arsenopyrite andpyrite, dissected by a dense net of quartz and carbonate veins. Pyrite forms
cubic, pentagon-dodecahedral, rarely cuboctahedral crystals.

0.2–0.25 **0.25–0.5 **0.5–1 **

33

Veinlet-disseminated ore type. Diamictites interspersed with pyrite, rarely
arsenopyrite, dissected by a dense net of quartz and carbonate veins fromhair-thick to 5 mm. Pyrite forms cubic, pentagon-dodecahedral, rarely
cuboctahedral crystals and their combined forms.

0.5–1
0.25–0.5
0.2–0.25

Vein ore type. Massive sulphide–quartz vein with a small number of fragmentsof diamictites, with nodules and inclusions of arsenopyrite, rarely pyrite. Cubic
pyrite crystals predominate.
1–2 **

0.25–0.5
0.14–0.25
0.07–0.14

Fraction(mm)

Streaky-vein ore type. Veins and veinlets of quartz and carbonate–quartz indiamictites, with inclusions of crystals and nodules of arsenopyrite, lessfrequently—pyrite. Pyrite forms cubic and pentagon-dodecahedral crystals.

PyriteCharacteristics of Samples for Selection of Monofractions

32

31

30

Nat-10M-161/10

26
27
28

29

TPM-1/1

Sample
No.*

23
24
25

FractionNo.

Table 1. Cont.

Minerals 2020, 10, 318

Minerals 2020, 10, 318

Table 2. The content of rhodium in the diﬀerently sized monofractions of arsenopyrite and pyrite
selected from vein and streaky-vein ores of the Natalkinskoe deposit.
Fraction No.

Sample No.

Monofraction (mm)Rh (ppm)

Arsenopyrite
123M-131/10

0.5–1
0.25–0.5
0.14–0.25

13
10
11

4
5
6
7

M-161/10

1–2
0.5–1
0.25–0.5
0.14–0.25

9.0
21
21
12

0.25–0.5
0.14–0.25
0.07–0.14

5.9
3.4
4.8

Pyrite8910TPM-1/1

Note: Ore sampling location and brief characteristics are presented in Table 1.

4.2. Results of the Study of Pt, Pd, Ru and Rh in Arsenopyrites and Pyrites by AAS-ADSSC
Data on the so-called “invisible” uniformly distributed impurity of Pt, Pd and Ru and the ratio of
their structural and surface-bound forms in arsenopyrites and pyrites of the Natalkinskoe deposit arepresented in Figures 5–7 and Tables 3–5. In all cases, highly deterministic (R2 = 0.82–1.0) dependencesof the average element content in the size sample on the speciﬁc surface area of the average crystal in it
were obtained. The highest R2 values were set for arsenopyrite: Pt, 0.92–0.99; Pd, 0.93–1.0 and Ru,
0.97–1.0 and were slightly lower for pyrite: Pt, 0.82–0.99; Pd, 0.94–0.99 and Ru, 0.91–0.99.

Figure 5. Dependences of the average content of evenly distributed platinum on the speciﬁc surface
area of an average crystal in size fractions of arsenopyrite (a) and pyrite (b) from vein, streaky-vein
and veinlet-disseminated ores of the Natalkinskoe deposit. The expressions for approximate curves
are shown here and in Figures 6–8, in which the pre-exponential factor (in bold) is an estimate of thecontent of the structural form of the element. The contents of structurally and surface-bound forms of
Pt, Pd, Ru and Rh in arsenopyrite and pyrite are given in Tables 3–6. Points 1–7 here and in Figures 6–8
correspond to the ore samples described in Table 1.

72

Minerals 2020, 10, 318

Figure 6. Dependences of the average content of evenly distributed palladium on the speciﬁc surface
area of an average crystal in size fractions of arsenopyrite (a) and pyrite (b) from vein, streaky-vein andveinlet-disseminated ores of the Natalkinskoe deposit.

Figure 7. Dependences of the average content of evenly distributed ruthenium on the speciﬁc surface
area of an average crystal in size fractions of arsenopyrite (a) and pyrite (b) from vein, streaky-vein andveinlet-disseminated ores of the Natalkinskoe deposit.

Figure 8. Dependences of the average content of evenly distributed rhodium on the speciﬁc surface
area of an average crystal in size fractions of arsenopyrite from vein and streaky-vein ores of the
Natalkinskoe deposit.

73

74

53–39

40–31

M-131/10

M-161/10

76–58

TPM-1/1

42–35

66–48

G-9/13M-129/10

Number of Crystals(Initial–Final
Sample)Sample No.
Interval ofMasses (mg)
0.12–0.18
0.19–0.28
0.29–0.69
0.79–1.32
1.34–3.64
0.10–0.14
0.15–0.20
0.21–0.22
0.23–0.23
0.24–0.28
0.29–0.33
0.34–0.46
0.10–0.12
0.13–0.24
0.27–0.37
0.48–0.73
1.01–1.45
0.10–0.17
0.18–0.32
0.34–0.52
0.54–1.05
1.07–2.15
0.40–0.62
0.66–1.01
1.03–1.39
1.42–4.31

Number of
Crystals

8
12
11
7
10

12
10
6
5
10
6
9
9
7
6
8
5
9
6
8
7
9
7
8
9
7

r (mm)

¯

0.58
0.9
1.2
2.58

0.14
0.24
0.43
0.88
1.58

0.11
0.19
0.31
0.59
1.3

0.12
0.17
0.22
0.23
0.26
0.31
0.41

0.15
0.24
0.55
1.14
2.34

0.456
0.529
0.582
0.751

0.284
0.339
0.412
0.524
0.637

0.262
0.314
0.371
0.459
0.597

0.271
0.304
0.33
0.336
0.35
0.371
0.406

0.292
0.339
0.448
0.572
0.727

Arsenopyritem (mg)

¯

2.151
1.866
1.694
1.312

3.456
2.873
2.369
1.872
1.541

3.745
3.114
2.664
2.143
1.645

3.672
3.262
2.97
2.945
2.827
2.664
2.412

3.411
2.873
2.189
1.722
1.355

Ssp (mm2 /mg)¯

Characteristics of the Final Sample

669

320

565

1067 ± 157
649 ± 136
436 ± 67236 ± 35
152 ± 281168 ± 278
838 ± 118532 ± 114
216 ± 75844

1488 ± 259
1052 ± 130
713 ± 221852 ± 98727 ± 115477 ± 156502 ± 77
2519 ± 443
1543 ± 282
802 ± 87438 ± 86202 ± 58

184

Ctot

¯

710 ± 109579 ± 56281 ± 54
100 ± 3055 ± 9C±Δ (ppm)

¯

499

288

536

728

145

¯

Cev

16

35

31

48

12

Cstr

Contents (ppm)

494

244

581

714

140

¯

Csur

Table 3. Results of platinum analysis of single crystal size selections of arsenopyrite and pyrite from vein, streaky-vein and veinlet-disseminated ores of the
Natalkinskoe deposit.

Minerals 2020, 10, 318

75

58–45

54–34

M-161/10

Nat-1078–50

UV-3/13

43–33

60–39

Nat-10

TPM-1/1

Number of Crystals(Initial–Final
Sample)Sample No.

0.10–0.12
0.13–0.16
0.17–0.21
0.22–0.31
0.34–1.13
0.30–0.55
0.60–1.96
2.24–2.71
2.84–4.14
4.48–16.2

12
10
8
10
5
8
8
7
6
5

0.11–0.24
0.25–0.35
0.36–0.51
0.52–0.75
0.76–1.10
1.19–2.16

11
8
8
9
7
7

0.10–0.12
0.13–0.18
0.21–0.33
0.41–1.65

0.21–0.41
0.42–0.56
0.58–0.98
1.17–2.31
3.73–14.33

8
8
9
8
6

8
9
6
10

Interval ofMasses (mg)Number of
Crystals¯¯r (mm)

0.45
1.46
2.51
3.65
11.5

0.11
0.15
0.19
0.27
0.65

0.11
0.16
0.27
0.67

0.18
0.31
0.41
0.67
0.99
1.7

0.3
0.49
0.74
1.65
9.21

Pyrite

0.448
0.663
0.795
0.9
1.32

0.28
0.311
0.336
0.378
0.507

0.28
0.317
0.378
0.512

0.33
0.396
0.434
0.512
0.583
0.698

0.366
0.431
0.495
0.646
1.147

Arsenopyritem (mg)

2.676
1.806
1.511
1.332
0.909

4.276
3.869
3.565
3.175
2.373

4.276
3.768
3.175
2.348

3.63
3.035
2.756
2.348
2.06
1.72

2.679
2.275
1.987
1.518
0.857

Ssp (mm2 /mg)¯

Characteristics of the Final Sample

Table 3. Cont.

574

990

56

1730 ± 1801320 ± 90
1200 ± 150
940 ± 57
380 ± 140239 ± 2197 ± 1756 ± 1131 ± 1213 ± 5252

688 ± 119375 ± 97375 ± 104210 ± 41
115 ± 3190 ± 221350 ± 100
940 ± 100610 ± 110272 ± 54

135

Ctot

¯

669 ± 85397 ± 102272 ± 51
125 ± 38
17 ± 14C±Δ (ppm)

¯

33

940

534

188

78

¯

Cev

3.7

68

41

14

4.2

Cstr

Contents (ppm)

35

890

462

190

89

¯

Csur

Minerals 2020, 10, 318

79–52

UV-3/13Interval ofMasses (mg)
0.10–0.18
0.20–0.29
0.31–0.39
0.41–0.58
0.60–1.00
1.05–3.60

Number of
Crystals

10
10
8
9
8
7

¯

0.16
0.24
0.35
0.52
0.78
2.69

m (mg)
Pyrite

¯

0.317
0.363
0.412
0.47
0.538
0.813

r (mm)

3.768
3.294
2.91
2.549
2.227
1.474

Ssp (mm2 /mg)¯

Characteristics of the Final Sample

716 ± 97372 ± 42360 ± 107316 ± 75342 ± 5484 ± 31

C±Δ (ppm)

¯

322

Ctot

¯

208

¯

Cev

37

Cstr

Contents (ppm)

189

¯

Csur

Ci mi ; Cev , average evenly distributed element content for all sizemi76

Number of Crystals(Initial–Final
Sample)

66–42

Sample No.

G-9/13

Interval ofMasses (mg)
0.12–0.18
0.19–0.28
0.29–0.69
0.79–1.32
1.34–3.64

Number of
Crystals
997
98

r (mm)

¯

0.15
0.25
0.57
1.05
2.42

0.292
0.345
0.453
0.556
0.735

Arsenopyritem (mg)

¯

3.411
2.857
2.16
1.766
1.339

Ssp (mm2 /mg)¯

Characteristics of the Final Sample

195 ± 43
97 ± 27
27 ± 626 ± 69.5 ± 3.5

C±Δ (ppm)

¯

40

Ctot

¯

27

¯

Cev

1.6

Cstr

Contents (ppm)

26

¯

Csur

Table 4. Results of palladium analysis of single crystal size selections of arsenopyrite and pyrite from vein, streaky-vein and veinlet-disseminated ores of the
Natalkinskoe deposit.

fractions; Cstr , structurally bound form content (C extrapolation to zero Ssp ); Csur , average content of surface-associated form.

fraction; C ± Δ, the average content of evenly distributed element ± one standard deviation; Ctot , average total content,

Note: Here and in Tables 4–6: m, the average mass of the crystal in size fraction; r, the average size of the crystal in size
fraction; Ssp , the speciﬁc surface area of average crystal in sizeNumber of Crystals(Initial–Final
Sample)Sample No.

Table 3. Cont.

Minerals 2020, 10, 318

77

40–24

60–39

M-161/10

Nat-1041–31

M-129/10

52–38

88–63

TPM-1/1

M-131/10

Number of Crystals(Initial–Final
Sample)Sample No.
Interval ofMasses (mg)
0.10–0.14
0.15–0.19
0.20–0.22
0.23–0.24
0.25–0.29
0.30–0.34
0.35–0.65
0.10–0.12
0.13–0.24
0.27–0.37
0.48–0.73
1.01–1.45
0.10–0.17
0.18–0.32
0.34–0.52
0.54–1.05
1.07–2.15
0.40–0.62
0.66–0.98
1.00–1.32
1.38–4.31
0.21–0.41
0.42–0.56
0.58–0.98
1.17–2.31
3.73–14.33

Number of
Crystals

10
11
8
10
9
7
8
9
6
5
6
5
8
6
8
8
8
8
5
6
5
9
8
8
7
7

¯

¯

r (mm)

0.27
0.51
0.77
1.64
8.42

0.58
0.85
1.12
2.87

0.14
0.26
0.43
0.87
1.58

0.11
0.2
0.31
0.6
1.25

0.12
0.17
0.21
0.23
0.27
0.32
0.44

0.353
0.438
0.501
0.646
1.113

0.456
0.518
0.569
0.777

0.284
0.35
0.412
0.523
0.637

0.262
0.321
0.371
0.461
0.59

0.271
0.304
0.324
0.336
0.353
0.373
0.416

Arsenopyritem (mg)

2.769
2.257
1.956
1.527
0.883

2.151
1.894
1.734
1.262

3.456
2.827
2.369
1.886
1.541

3.745
3.091
2.664
2.125
1.671

3.672
3.262
3
2.945
2.769
2.609
2.36

Ssp (mm2 /mg)¯

Characteristics of the Final Sample

Table 4. Cont.

32

39

12

83 ± 18
46 ± 18
34 ± 6
11 ± 796 ± 1833 ± 918 ± 59.8 ± 4.2
1.9 ± 0.897361 ± 39
221 ± 47133 ± 46
54 ± 19
20 ± 4122 ± 3177 ± 1052 ± 1025 ± 6
10 ± 2

142

Ctot

¯

245 ± 69
196 ± 45132 ± 30
117 ± 27
101 ± 2375 ± 2073 ± 18

C±Δ (ppm)

¯

7.7

31

30

83

114

¯

Cev

0.4

0.6

2

2.6

5.9

Cstr

Contents (ppm)

8.1

31

27

82

114

¯

Csur

Minerals 2020, 10, 318

78

60–41

79–56

Nat-10UV-3/13

43–33

TPM-1/1

54–44

76–58

UV-3/13

M-161/10

Number of Crystals(Initial–Final
Sample)Sample No.

0.29–0.51
0.55–1.63
1.72–2.57
2.66–4.14
4.48–16.2
0.10–0.18
0.20–0.29
0.31–0.39
0.41–0.58
0.60–1.00
1.05–3.60

8
7
8
9
9

10
9
11
12
9
5

0.10–0.12
0.13–0.18
0.21–0.33
0.41–1.65

10869
0.10–0.12
0.13–0.17
0.18–0.22
0.24–0.31
0.34–1.13

0.11–0.24
0.25–0.35
0.35–0.51
0.52–0.75
0.76–1.10
1.19–2.16

12
10
9
10
8
9

10
9
7
9
9

Interval ofMasses (mg)Number of
Crystals¯r (mm)

0.11
0.15
0.2
0.28
0.54

0.11
0.16
0.29
0.7

0.17
0.25
0.35
0.51
0.83
3.2

Pyrite

0.324
0.368
0.412
0.467
0.55
0.862

0.438
0.628
0.757
0.864
1.287

0.28
0.311
0.342
0.383
0.476

0.28
0.317
0.378
0.519

0.324
0.396
0.434
0.514
0.583
0.685

Arsenopyrite
0.17
0.31
0.41
0.68
0.99
1.61

0.42
1.24
2.17
3.23
10.66

¯

m (mg)

3.705
3.25
2.91
2.566
2.187
1.393

2.741
1.908
1.584
1.387
0.932

4.276
3.869
3.509
3.143
2.518

4.276
3.768
3.099
2.309

3.705
3.035
2.756
2.331
2.06
1.749

Ssp (mm2 /mg)¯

Characteristics of the Final Sample

Table 4. Cont.

73

8.7

22

35 ± 614 ± 311 ± 37.9 ± 0.8
2.7 ± 0.577 ± 1145 ± 1237 ± 425 ± 218 ± 34.2 ± 1.2

58

21

Ctot

¯

142 ± 11111 ± 880 ± 662 ± 337 ± 7156 ± 1797 ± 1158 ± 1225 ± 582 ± 1144 ± 535 ± 617 ± 310 ± 26.7 ± 1.1

C±Δ (ppm)

¯

15

6.3

67

53

18

¯

Cev

1

1.1

5.2

3.2

0.8

Cstr

Contents (ppm)

18

4.9

62

50

25

¯

Csur

Minerals 2020, 10, 318

79

40–29

40–33

M-131/10

M-161/10

85–61

TPM-1/1

43–33

64–42

G-9/13M-129/10

Number of Crystals(Initial–Final Sample)

Sample No.
Interval ofMasses (mg)
0.12–0.20
0.22–0.33
0.40–0.69
0.79–1.34
1.39–3.64
0.10–0.14
0.15–0.19
0.20–0.22
0.23–0.24
0.25–0.29
0.30–0.34
0.35–0.65
0.10–0.12
0.13–0.24
0.27–0.37
0.48–0.73
1.01–1.45
0.10–0.17
0.18–0.32
0.34–0.52
0.54–1.05
1.07–2.15
0.40–0.62
0.66–1.01
1.03–1.39
1.42–4.31

Number of
Crystals
12
9
6
7
8
9
12
10
9
8
6
7
9
6
5
7
6
6
5
7
6
5
9
8
9
7

r (mm)

¯

0.56
0.91
1.2
2.62

0.15
0.26
0.44
0.9
1.7

0.11
0.2
0.3
0.61
1.27

0.12
0.17
0.21
0.23
0.28
0.32
0.45

0.16
0.27
0.57
1.1
2.49

0.451
0.53
0.582
0.755

0.292
0.35
0.416
0.529
0.65

0.262
0.321
0.366
0.464
0.592

0.271
0.304
0.324
0.336
0.358
0.373
0.42

0.296
0.353
0.453
0.565
0.742

Arsenopyritem (mg)

¯

2.179
1.852
1.694
1.305

3.411
2.827
2.36
1.866
1.505

3.745
3.091
2.679
2.118
1.656

3.672
3.262
3
2.945
2.746
2.609
2.352

3.286
2.769
2.16
1.741
1.327

Ssp (mm2 /mg)¯

Characteristics of the Final Sample

656

179

331

789 ± 92432 ± 121263 ± 33
129 ± 3063 ± 15667 ± 107422 ± 75309 ± 42166 ± 33

1318

2424 ± 512
1678 ± 316
1321 ± 159
1108 ± 225
904 ± 197
832 ± 298
530 ± 862269 ± 285
1179 ± 196
812 ± 63516 ± 63
231 ± 59

283

Ctot

¯

1034 ± 177
547 ± 76339 ± 149
160 ± 53
82 ± 26

C±Δ (ppm)

¯

297

165

537

1129

193

¯

Cev

20

10

48

41

18

Cstr

Contents (ppm)

290

162

520

1083

190

¯

Csur

Table 5. Results of ruthenium analysis of single crystal size selections of arsenopyrite and pyrite from vein, streaky-vein and veinlet-disseminated ores of the
Natalkinskoe deposit.

Minerals 2020, 10, 318

80

58–43

58–40

M-161/10

Nat-10

78–58

UV-3/13

43–29

59–35

Nat-10

TPM-1/1

Number of Crystals(Initial–Final Sample)

Sample No.

0.10–0.12
0.13–0.16
0.17–0.21
0.22–0.31
0.34–1.13
0.29–0.55
0.60–1.72
1.81–2.66
2.71–4.48
5.78–16.2

10
10
6
10
7
6
10
7
9
8

0.11–0.24
0.25–0.35
0.36–0.51
0.52–0.75
0.76–1.10
1.19–2.16

10
9
10
10
9
10

0.10–0.12
0.13–0.18
0.21–0.33
0.41–1.65

0.21–0.41
0.42–0.56
0.58–0.98
1.17–2.31
3.73–14.33

5
8
8
8
6

8
8
6
7

Interval ofMasses (mg)Number of
Crystals¯¯r (mm)

0.42
1.23
2.32
3.58
9.91

0.11
0.15
0.19
0.27
0.58

0.11
0.16
0.27
0.73

0.18
0.32
0.4
0.66
0.98
1.63

0.30
0.50
0.80
1.65
9.07

Pyrite

0.438
0.627
0.774
0.895
1.256

0.28
0.311
0.336
0.378
0.488

0.28
0.317
0.378
0.527

0.33
0.4
0.431
0.509
0.581
0.688

0.336
0.434
0.508
0.646
1.141

Arsenopyritem (mg)

2.741
1.918
1.549
1.343
0.955

4.276
3.869
3.565
3.175
2.464

4.276
3.768
3.175
2.283

3.63
3
2.787
2.355
2.067
1.742

2.679
2.260
1.936
1.518
0.861

Ssp (mm2 /mg)¯

Characteristics of the Final Sample

Table 5. Cont.

686

850

48

1560 ± 350
1320 ± 130
930 ± 290
540 ± 70330 ± 90189 ± 86
109 ± 2258 ± 1533 ± 514 ± 4329

941 ± 172428 ± 69372 ± 65245 ± 49
211 ± 56
103 ± 201720 ± 300
1280 ± 210
605 ± 160
235 ± 90

128

Ctot

¯

729 ± 262
375 ± 103
187 ± 51
124 ± 49
24 ± 10C±Δ (ppm)

¯

35

700

574

238

79

¯

Cev

4.9

34

23

19

5.7

Cstr

Contents (ppm)

30

690

577

218

79

¯

Csur

Minerals 2020, 10, 318

79–55

UV-3/13

0.10–0.18
0.20–0.29
0.31–0.39
0.41–0.58
0.60–1.00
1.05–3.60

10
11
10
9
9
6

¯

0.16
0.24
0.36
0.52
0.83
2.9

Pyritem (mg)

¯

0.317
0.363
0.416
0.47
0.558
0.834

r (mm)

3.768
3.294
2.884
2.549
2.187
1.439

Ssp (mm2 /mg)¯

Characteristics of the Final Sample
Interval ofMasses (mg)Number of
Crystals
1840 ± 400
1085 ± 149
577 ± 129384 ± 74248 ± 3967 ± 23

C±Δ (ppm)

¯

454

Ctot

¯

272

¯

Cev

9.9

Cstr

Contents (ppm)

Number of Crystals(Initial–Final Sample)

43–31

38–28

40–28

Sample No.

M-129/10

M-131/10

M-161/10

Interval ofMasses (mg)
0.10–0.12
0.13–0.24
0.27–0.37
0.48–0.73
1.01–1.45
0.10–0.17
0.18–0.31
0.34–0.52
0.54–1.04
1.07–2.15
0.40–0.62
0.66–1.01
1.03–1.39
1.42–4.31

Number of
Crystals
9
5
6
6
5

81

5
4
4
7
8
8
7
7
6

0.54
0.93
1.26
2.65

0.14
0.26
0.49
0.83
1.64

0.11
0.21
0.31
0.60
1.29

m (mg)

¯

0.446
0.534
0.592
0.757

0.284
0.350
0.431
0.514
0.650

0.262
0.324
0.371
0.461
0.595

r (mm)

¯

2.210
1.840
1.669
1.297

3.456
2.827
2.275
1.910
1.527

3.745
3.000
2.664
2.125
1.647

Ssp (mm2 /mg)¯

Characteristics of the Final Sample
Ctot

801

211

360

3108 ± 491
1473 ± 394
973 ± 159
483 ± 78296 ± 85653 ± 391
453 ± 130368 ± 146
224 ± 34
103 ± 19782 ± 155387 ± 55
320 ± 89
158 ± 46¯

C±Δ (ppm)

¯

309

190

715

¯

Cev

17

36

45

Cstr

Contents (ppm)

Table 6. Results of rhodium analysis of size selections of arsenopyrite single crystals from vein and streaky-vein ores of the Natalkinskoe deposit.

Number of Crystals(Initial–Final Sample)

Sample No.

Table 5. Cont.

299

154

678

¯

Csur

330

¯

Csur

Minerals 2020, 10, 318

Minerals 2020, 10, 318

Extrapolation of these dependences to zero speciﬁc surface area (Ssp , mm2 /mg, Tables 3–5), i.e.,
on a conditionally inﬁnite crystal in the ﬁrst case of arsenopyrite, and in the second pyrite, yields thefollowing estimates of the contents of the structural form of the elements (Cstr , ppm): for arsenopyrite
Pt, 4.2–48; Pd, 0.4–5.9 and Ru, 5.7–48; for pyrite Pt, 3.7–68; Pd, 1.0–5.2 and Ru 4.9–34 (Figures 5–7).
The average contents of surface-bound form of elements (Csur , ppm) are given in Tables 3–5 and forarsenopyrite amount to the following values: Pt, 89–714; Pd, 8.1–114 and Ru, 79–1083; for pyrite Pt,
35–890; Pd, 4.9–62 and Ru, 30–690.
By now, the range of PGE studied by the AAS-ADSSC method has expanded. The ﬁrst results forthe forms of occurrence of Rh, one more element of the platinum group in crystals of arsenopyrite,
were obtained. These data are still limited and preliminary. The results of the study of the content of
evenly distributed Rh and the ratio of its structural and surface-bound forms in arsenopyrites fromvein and streaky-vein ores of the Natalkinskoe deposit are presented in Figure 8 and in Table 6. In allcases, as well as for Pt, Pd and Ru, highly deterministic (R2 = 0.89–1.0) dependences of the average
Rh content in the size fraction on the speciﬁc surface area of the average arsenopyrite crystal in it
were obtained. Extrapolation of these dependences to a zero-speciﬁc surface, i.e., to a conditionallyinﬁnite crystal, yielded estimates of the contents of the structural form of this element of 17–45 ppm.
The average grades of the surface-associated forms were 154–678 ppm for arsenopyrite (see Table 6).
The data presented (PCA and AAS-ADSSC) suggest that arsenopyrite also acts as a concentrator
mineral for Rh. Structural and surface-bound forms in arsenopyrite (as well as in pyrite), as the main
uniformly distributed forms, are characteristic not only of Pt, Pd and Ru, but also Rh. The data obtained
conﬁrms that arsenopyrites and pyrites are major mineral concentrators of PGE at the Natalkinskoe
gold ore deposit. It can also be assumed that the size dependence of the content of evenly distributed
Pt, Pd, Ru and Rh, and the conﬁnement of a signiﬁcant part to the surface of the crystals, is not aunique phenomenon for sulphide minerals of ore deposits. Structurally and superﬁcially bound forms,
as the main uniformly distributed forms, may be typical of both the remaining PGE (Os and Ir) and ofother noble metals—Au and Ag [11,37–39,41–45].
New data for Pt and Pd forms in arsenopyrites, both structural and surface-bound, conﬁrmed
earlier conclusions [11]. The uniformly distributed Pt, Pd and Ru and, as we believe, Rh in arsenopyrites
and pyrites are forms of elements which are chemically bound in the structure of these minerals and innanoscale nonautonomous phases (NAPs) on the surface of arsenopyrite and pyrite crystals. The latter
usually dominates over the former according to the content of these elements. Recent experimental
and natural data [37,38,46] suggest that Pt, Pd, Ru and Rh, as well as Au, can be incorporated into NAP
formed on the surface of arsenopyrite and pyrite crystals in a very thin surface layer (~100–500 nm).
These elements are absorbed by NAP from the hydrothermal solution even more eﬃciently than gold,
which is incompatible in pyrite.
There is currently high demand for studies of the surface layer of sulphide minerals as possible
hosts for “invisible” forms of trace elements, the nature of their distribution and concentration levels.
Nanoscale NAPs enriched with Pt, Pd, Ru and Rh form a part of the so-called impurity component
on the surface of arsenopyrite and pyrite crystals. The importance of its investigation is obvious.
It is assumed that there is a common mechanism of absorption of all noble metals for the studied
sulphides due to the active role of the surface during the growth of the crystal from a supersaturatedhydrothermal solution [38,39]. This mechanism is closely related to the conditions of ore formation.
This is an important point that must be taken into account in practical terms, for example, in theconcentration of ores that include sulphides, in our case arsenopyrite and pyrite.
The structurally bound form of PGE, primarily Pt, despite a lower content compared to the surface
form, is also extremely important. As shown earlier, data on the structural form of Au, Ag and Pt
make it possible to make comparative estimates of the contents of these elements in the ore-forming
ﬂuids that form gold deposits. It is the only form of the element occurrence that can be used as anindicator of element activity in hydrothermal solution [37,46]. For example, the experimental data on
the distribution and segregation of trace elements in crystals of ore minerals grown in hydrothermal82

Minerals 2020, 10, 318

systems were used to assess the content of Au and Pt in the ore-forming ﬂuid (Caq ) of the Natalkinskoe
and Degdekan gold deposits. The content of the element in the structural form in natural pyrite,
and the distribution coeﬃcient (Dstr ) for the same form between pyrite and hydrothermal solution
according to experimental data, were used for estimating element content in the ore-forming ﬂuid [38].
In Table 7 these results are supplemented by the data from the present work on samples M-161/10,
TPM-1/1 and Nat-10.
Table 7. Evaluation of gold and platinum grade in ore-forming ﬂuids by pyrite composition using
experimental distribution coeﬃcients Dstr .
Number of
Crystals(Initial–Final
Sample)Element

Cstr (ppm)

Csur (ppm)

Caq (ppm)

Au/Ptaq

95–59
92–71

Au
Pt

0.2
2.0

0.5
7.1

2.1
0.1

21

DG-10/14

67–44
34–22

Au
Pt

0.1
5.1

1.1
102

1.3
0.2

6.5

UV-3/13

80–52
79–52

Au
Pt

0.3
37

1.4
189

2.9
1.8

1.6

M-161/10

58–40
58–45

Au
Pt

0.2
68

3.6
890

1.8
3.2

0.6

TPM-1/1

43–30
43–33

Au
Pt

0.3
41

3.4
462

3.2
2.0

1.6

Nat-10

60–39
54–34

Au
Pt

1.8
3.6

2.7
35

18
0.2

90

Sample No.

Deposit

M-163/10
Degdekan

Natalkinskoe¯

Note: Cstr , structurally bound element content; Csur , superﬁcially bound element content; Caq , element content in
the ore-forming ﬂuid.

The average value of Dstr is 0.1 for Au and 21 for Pt [41,46]. The evaluation of Au and Pt content
in the ore-forming ﬂuid showed that Au content in the ﬂuids of the Degdekan and Natalkinskoe gold
deposits was mainly at the level of 1.3–3.2 ppm, and it reached a maximum value of 18 ppm for the
Nat-10 sample. Pt grade ranged from 0.1 to 3.2 ppm and, on average, was clearly elevated at the
Natalkinskoe deposit compared to Degdekan. The ﬂuid of the Natalkinskoe deposit was rich in bothelements compared to the ﬂuid of the Degdekan deposit. Au and Pt contents in solution were also
comparable for most samples from the Natalkinskoe deposit (with the exception of the anomalous Au
in sample Nat-10), whereas Au dominated over Pt in the Degdekan deposit. This is another argument
in favor of studying platinum-bearing gold deposits, of which the Natalkinskoe deposit is no exception.
Information about the Pt content in the ﬂuid is of particular value in view of the complex behavior ofPt and the heterogeneity of its distribution under the capture of gas–liquid inclusions, the compositionof which is determined by LA-ICP-MS as a proxy of the ore ﬂuid composition [47].
4.3. Study of the Chemical Composition of the Surface of Crystals of Arsenopyrite and Pyrite Using theSEM-EDX Method
Nearly isometric pseudo-orthorhombic crystals of arsenopyrite and cubic crystals of pyrite without
visible inclusions of ore minerals and with a minimum number of non-metallic microinclusions (mainly
quartz and carbonate) were selected for the study (Figure 9). The so-called “clean” areas virtuallylacking visible inclusions and signiﬁcant defects were then selected to study the composition of the“invisible” impurity component on the surface of these crystals (Figure 10). The size (in μm) of suchareas was 10 × 10, 20 × 20, and it was at least 30 × 30. Sections with low roughness were selected using
the scanning probe microscopy technique (see above).

83Minerals 2020, 10, 318

Figure 9. Crystals of arsenopyrite (Asp) and pyrite (Py) with microinclusions of carbonate (Cb) andquartz (Qtz). Surface areas of arsenopyrite (1) and pyrite (2) crystals, where measurements were carriedout. Their enlarged fragments are shown in Figure 10.

Figure 10. Enlarged fragments of arsenopyrite (Asp) and pyrite (Py) crystals with “clean” areas, where
measurements were carried out (see Table 8, sample М-129/10, N 3, n 3 and Table 9, sample М-129/10,
N 1, n 2). The bottom of the Figure 9 shows the X-ray energy spectra of arsenopyrite (see Table 8,
N 3, n 3).

84Minerals 2020, 10, 318

Faces with striation of the same crystallographic type were selected during scanning in order toeliminate the eﬀect of capturing various impurities by faces of diﬀerent crystallographic types during
crystal growth: for arsenopyrite—prism {141}, for pyrite—cube {100}.
The surfaces of arsenopyrites and pyrites selected from the Natalkinskoe ore deposit (samples
G-9/13, TPM-1/1, M-129/10, M-131/10, M-161/10 and Nat-10) clearly demonstrated the peculiarities of
its elemental composition. On the “clean” surface of arsenopyrite, PGEs were thus represented by thefollowing elements (in wt%): Pt (up to 2.1), Ru (1.4), Ir (up to 0.7) and Os (up to 0.7). Other ore elements
were also determined (in wt%): U (up to 2.0), Hg (up to 1.1), Au (up to 1.0), rarely Ag (up to 0.7) andCu (up to 0.6) (Table 8). The only PGE present on the “clean” surface of pyrite was Pt (up to 2.9 wt%),
and other ore elements were represented by As (up to 2.2 wt%) and Cu (up to 0.7 wt%) (Table 9).
The SEM-EDX method conﬁrmed the presence of Au and Pt on the surface of arsenopyrite crystals
from the Natalkinskoe deposit, and high contents of them were detected earlier [11]. It is the ﬁrst timesuch PGEs as Ru, Os, Ir and a number of ore elements—Ag, Hg, Cu and U—were detected on thesurface of arsenopyrite. Pt and As are shown for the ﬁrst time to be constantly present on the surfaceof pyrite crystals. Comparing the data from the AAS-ADSSC and SEM-EDX methods regarding thesurface contents of Au, Pt and Ru, it is necessary to recall that the data on Csur of AAS-ADSSC method
are related to the entire average crystal, whereas the SEM-EDX method virtually deﬁnes local content of
the element at the local area (“point”) on the surface. It would be more correct to compare these resultswith the content of the element within NAP, which can be calculated under certain assumptions [46].
According to such calculations, Au in NAP on pyrite amounts to 0.3–0.5 wt% [38], which is comparable
by an order of magnitude to the data in Tables 8 and 9 on Au and Pt. Increased platinum on pyrite(see Table 9) is qualitatively consistent with the fact that, according to AAS-ADSSC, Csur Pt > Csur Au(see Table 7).

85

86

32.6
32.1
32.4

34.0
34.7

30.8
31.9

33.3
33.2
33.1

1
2
3

1
2

1
2

1
2
3

3

4

6

5

30.0
28.9

1

1
2

1
2
3

2

Fe

31.6
32.4
31.9

n

N
As

42.4
42.9
43.2

41.2
41.7

40.2
41.5

41.8
43.4
43.1

36.3
39.0

36.0
35.6
36.8

S

19.0
19.1
19.1

17.1
19.0

15.1
13.8

16.1
14.2
15.4

20.4
20.6

23.4
20.2
20.8

C

1.0
0.6
0.4

3.9
1.9

4.6
4.8

3.9
2.3
1.1

4.3
4.6

2.8
3.4
3.7

O

1.1
0.8
0.5

3.2
2.1

2.7
1.8

1.5
1.4
1.0

2.5
2.5

2.0
2.5
1.5

Al

0.6
0.5
0.5

1.0
1.4

-

-

1.2
0.5

1.2
1.3
0.6

Si

0.2
0.2
0.1

0.1
0.1

-

-

0.1
0.2

0.2
0.2
0.1

-

-

0.4
0.4

1.0
0.9
0.9

1.1
0.9

0.7
1.1
1.0

Hg

-

-

0.3
0.4

0.7
1.0
0.9

0.9
0.8

0.7
0.9
0.9

Au

0.3
0.1
0.6

0.1

-

0.2
0.7

0.1
0.2

0.2
0.3
0.3

Ag

0.2
0.2
0.6

0.6
0.5

-

-

-

-

Cu

0.2
0.1
0.1

0.2

-

-

-

-

Ti

Pt

0.6
1.2
0.8

2.1
1.1

0.4
0.4

0.5
0.9
0.9

1.1
0.4

0.4
0.6
0.8

0.2
0.2
0.2

-

0.2
0.2

0.6
0.6
0.7

-

-

Os

0.3
0.3
0.3

-

0.3
0.2

0.4
0.7
0.7

-

-

Ir

0.3
0.3
0.2

-

1.4
1.2

1.0
1.0

-

-

Ru

U

0.3
0.3
0.3

-

0.4
0.6

0.9
1.3
1.2

2.0
1.4

0.8
1.5
1.6

Note: Pd, Rh, F, Cl, Pb, Zn, Mo, W, Bi and REE were not detected. Analytical lines: FeKα , AsKα , SKα , CKα , OKα , AlKα , SiKα , HgLα , AuLα , AgLα , CuKα , TiKα , PtLα , OsLα , IrLα , RuLα
and ULα . Here and in Table 9: N—the serial number of the crystal, n—the serial number of the measurement area on the crystal surface; hyphen—element is not detected; italics—theelements content below the detection limit (<0.5 wt%), so their presence is problematic but without taking them into account the sum total does not reach 100%; the main composition isin bold.

M-161/10

M-129/10

Sample No.

Table 8. Chemical composition (in wt%) of arsenopyrite crystal surfaces with micron-size rare inclusions of non-metallic minerals. Samples from vein ores of the
Natalkinskoe deposit.

Minerals 2020, 10, 318

Minerals 2020, 10, 318

Table 9. Chemical composition (in wt%) of pyrite crystal surfaces with micron-size rare inclusions of
non-metallic minerals. Samples from vein ores of the Natalkinskoe deposit.
Sample No.

N

n

Fe

1

1
2
3

2

Al

Si

Ca

Na

As

Cu

Ti

Pt

38.5 43.1 5.3
36.9 39.2 7.0
41.0 47.9 2.97.7 1.2
10.7 1.3
2.6 1.4

0.2
0.3
0.2

-

-

1.3
1.2
1.7

0.7
0.7
0.5

0.3
0.4
0.2

1.7
2.3
1.6

1
2
3

39.8 44.1 5.1
35.9 42.7 5.6
38.8 47.0 3.3

6.8 1.1
10.2 0.8
6.2 0.9

0.2
1.1
0.9

-

-

1.0
1.0
1.1

0.6
0.3

0.2
0.1
0.1

1.7
2.0
1.4

3

1
2

37.0 38.3 4.3
39.5 48.4 2.3

13.2 0.9
3.0 0.8

0.7
0.1

0.2
0.1

0.3
0.2

1.8
2.2

0.6
0.5

-

2.7
2.9

4

1
2

37.4 37.5 3.1
40.5 46.5 3.6

16.4 1.3
5.6 1.4

1.0
0.1

-

0.2

1.1
1.4

0.4
0.2

0.1
-

1.7
0.5

5

1
2

41.7 47.2 2.6
40.7 47.5 3.0

3.3
4.0

0.1
0.2

-

-

2.0
2.0

0.3
0.2

-

1.9
1.5

M-129/10

M-161/10

S

C

O

0.9
0.9

Note: Au, Ag, Hg, Pd, Os, Ir, Ru, Rh, U, F, Cl, Pb, Zn, Mo, W, Bi and REE were not detected. Analytical lines: FeKα ,SKα , CKα , OKα , AlKα , SiKα , CaKα , NaKα , AsKα , CuKα , TiKα and PtLα .

4.4. Study of the Chemical Composition of the Surface of Arsenopyrite and Pyrite Crystals by LA-ICP-MS
When studying the surface layers of arsenopyrite and pyrite crystals using the LA-ICP-MS method,
attention was largely focused on Pt and Pd. Ten arsenopyrite crystals and six pyrite crystals wereanalyzed (Tables 10 and 11). For all crystals of arsenopyrite (samples M-129/10, M-131/10, M-161/10
and TPM-1/1) and pyrite (samples M-161/10 and TPM-1/1) a track depth of 0.5–0.6 μm per laser pass
was obtained. Unfortunately, we do not know the NAP thickness of the studied crystals. If we focus on
the previously obtained data [44,48], the thickness of NAP on pyrite is ≤0.5 μm, and for arsenopyrite it
is, according to preliminary data, much less [11]. The ﬁrst pass of the laser thus best represents thecomposition of NAP, although signiﬁcant “contamination” by the material underlying NAP is likely.
The parameters of the laser were adjusted for one of the arsenopyrite crystals selected for detailed
study (see Table 10, sample M-129/10, crystal 2), in such a way that a layer of ~100 nm thick would beburned in one pass (see above). In this case, the maximum depth of the laser-dig groove was 600 nm.
Figure 11 shows graphs of the detailed distribution of Pt and Pd contents in the surface layer of Crystal
2 from the M-129/10 sample.

Figure 11. Distribution of the average content (C) of platinum and palladium in the surface layerof arsenopyrite crystal (sample M-129/10, Crystal 2) according to the results of LA-ICP-MS analysis(the Natalkinskoe deposit).

87Minerals 2020, 10, 318

Table 10. Results of LA-ICP-MS analysis of the surface of arsenopyrite crystals for Pt and Pd(the Natalkinskoe deposit).
Sample No.

Crystal No. (Size, mm)

Track No.

1 (1.0)

Pt

Pd

1
2

2.9
1.4

1.8
1.8

2 (1.1)

1

3.1
2.8
1.4
-

2.5
1.0
1.8
1.8
1.0

1 (1.3)

1
2

0.4
0.4

0.7
0.5

2 (1.3)

1
2

-

0.1
0.2

3 (1.0)

1
2

<0.4
-

0.1
0.1

1 (1.0)

1
2
3
4
5

0.4
23
0.4
<0.4
6.2

0.4
3.8
0.2
0.1
1.4

2 (1.2)

1
2

<0.4
-

<0.1
0.6

1 (1.3)

1

0.4

0.8

2 (1.5)

1

0.4

1.2

3 (1.5)

1
2

<0.4
0.4

0.9
4.8

M-129/10

M-131/10

M-161/10

TPM-1/1

Contents of Elements (ppm)

Note: Here and in Tables 11–13:hyphen—element not detected; sign “<”—below detection limit. Bold—results of adetailed study of Pt and Pd contents to the depth of 600 nm with a step of 100 nm (see Figure 11).

Table 11. Results of LA-ICP-MS analysis of the surface of pyrite crystals for Pt and Pd(the Natalkinskoe deposit).
Sample No.

Track No.

Pt

Pd

1 (1.0)

1
2
3
4

1.6
0.4
-

5.9
0.1
0.2
0.1

2 (1.2)

1
2
3

<0.4
<0.4
<0.4

0.9
0.2
0.4

3 (1.5)

1
2

-

0.2
0.4

1 (1.8)

1
2

0.4
-

3.8
0.1

2 (1.7)

1
2

-

0.7
0.1

3 (1.9)

1
2

0.4
-

0.1
0.7

M-161/10

TPM-1/1

Contents of Elements (ppm)

Crystal No. (Size, mm)88Minerals 2020, 10, 318

A generalization of the results of previous studies [11] and the new data suggest that the increased
content of Pt and Pd is conﬁned to the surface layer of not only arsenopyrite crystals, but also pyrite.
It can also be assumed that the tendency for a Pt and Pd concentration in the surﬁcial layer at <500 nm
thick is common for both minerals (see Figure 11 and Table 11). Several exceptions can be noted inTable 10, however, which may be due to inhomogeneity in the impurity distribution because of theevolution of NAP during crystal growth [39].
Determination of the contents of a wider range of elements belonging to the group of PGEs—Pt,
Pd, Rh and Ir—was carried out in the example of studying the surfaces of arsenopyrite and pyrite
crystals taken from the UV-3/13 sample (Figures 12 and 13). The surface layer of these crystals wasstudied to a depth of 6 μm (depth per pass is 1 μm). Ru and Os contents were not determined dueto their absence in the standards we used. According to LA-ICP-MS analysis, the surface layers ofarsenopyrite and pyrite, along with Pt and Pd, concentrated PGEs such as Rh and Ir. The contents ofelements (ppm) in the surface layers of arsenopyrite were Pt up to 5.2, Pd up to 8.4, Rh up to 8.9 andIr up to 0.2, and in the surface layers of pyrite were Pt up to 13, Pd up to 896, Rh up to 6.7 and Ir up
to 4.1. The maximum contents of Pt, Pd and Ir were established in pyrite, and Rh was established inarsenopyrite. Like Pt and Pd, all the highest contents of Rh and Ir were conﬁned to the ﬁrst removed
layer (1 μm thick). There was a distinct tendency for the content of all these elements to decrease withdepth (Tables 12 and 13).
High Pd content and Pd/Pt ratios on the pyrite surface (see Table 13) contradict the AAS-ADSSC
and SEM-EDX data presented above, according to which the Pt content was more than an order ofmagnitude higher than Pd. The reason for this contradiction requires special research and may berelated to the existence of an unidentiﬁed surface form of Pd (for example, Pd nanoparticles), which isinhomogeneously distributed over crystals. The SEM-EDX method was also local (locality is evenhigher than for LA-ICP-MS), and therefore cannot serve to conﬁrm this hypothesis. In the AAS-ADSSC
method, crystals with such an excess of Pd over the average content simply will not be included in theﬁnal sample.

Figure 12. Arsenopyrite Crystals 1 and 2 (sample UV-3/13) from ores of the Natalkinskoe deposit.
Tracks 1–9, numbers of laser passes over the crystal surface. The results of LA-ICP-MS analysis arepresented in Table 12.

89Minerals 2020, 10, 318

Figure 13. Pyrite Crystals 1 and 2 (sample UV-3/13) from ores of the Natalkinskoe deposit. Tracks 1–8,
numbers of laser passes over the crystal surface. The results of LA-ICP-MS analysis are presented inTable 13.
Table 12. Results of LA-ICP-MS analysis (in ppm) of the surface of arsenopyrite crystals (1 and 2) for
Pt, Pd, Rh and Ir (sample UV-3/13) (the Natalkinskoe deposit).
n

Pt

Pd

Rh

Ir

Pt

Pd

Rh

Ir

Pt

Pd

Rh

Ir

2.3
0.5
1.6
0.4
0.3
-

0.1
<0.1
-

7.5
0.7
0.5
0.2
2.1
0.2

-

8.9
3.6
3.7
1.1
1.5
0.7

<0.1

Crystal 1
Track 1
1
2
3
4
5
6

0.7
-

1.6
<0.2
0.3
0.4
0.2
<0.2

7.6
2.5
0.2
<0.2
0.3
0.2

Track 2
0.2
<0.1
-

<0.7
0.9
-

5.2
2.2
7.6
3.2
2.0
2.0

1.6
1.0
0.3
0.5
0.4
0.6

Track 3
<0.1
<0.1
-

<0.7
-

7.9
3.4
3.5
2.9
1.4
1.5

Crystal 2
Track 4
1
2
3
4
5
6

<0.7
-

0.8
0.2
0.2
0.3
<0.2
-

1
2
3
4
5
6

1.3
<0.7
-

8.2
0.6
<0.2
0.4
<0.2
0.4

1.2
0.5
0.2
0.5
<0.2
<0.2

Track 5
0.1
0.1
-

5.2
-

2.9
1.2
<0.2
1.4
0.4
0.9

-

<0.7
<0.7
-

2.8
1.5
1.8
<0.2
0.2
<0.2

Track 7
1.9
1.7
0.8
0.9
0.4
<0.2

0.6
0.2
0.2
0.7
0.4
<0.2

Track 6
-

0.7
-

5.8
4.3
1.7
1.5
1.3
0.7

<0.1
-

1.5
-

8.4
3.7
2.5
3.7
3.6
2.7

Track 8
6.4
0.5
0.8
0.3
0.9
-

Track 9Note: Here and in Table 13: n—numbers of laser passes along the track (depth per pass is 1.0 μm).

90

Minerals 2020, 10, 318

Table 13. Results of LA-ICP-MS analysis (in ppm) of the surface of pyrite crystals (1 and 2) for Pt, Pd,
Rh and Ir (sample UV-3/13) (the Natalkinskoe deposit).
n

Pt

Pd

Rh

Ir

Pt

Pd

Rh

Ir

Pt

Pd

Rh

Ir

Pt

Pd

1.1
0.5
0.2
<0.2
-

2.2
2.2
1.1
<0.7
<0.7
-

70
109
47
63
31
34

4.1
1.9
1.5
0.8
0.5
0.5

7.7
5.3
8.3
0.9
2.6
2.2

693
715
461
301
135
143

Rh

Ir

Crystal 1
Track 1
1
2
3
4
5
6

2.5
1.0
0.5
0.3
-

118
54
26
10
7.7
2.6

0.9
0.7
1.8
-

Track 2
0.5
0.3
0.1
-

3.6
1.1
<0.7
0.9
0.7
0.9

135
108
54
56
28
39

1.6
<0.2
0.3
<0.2
<0.2
<0.2

Track 3
0.4
0.4
0.6
<0.2
-

1.8
1.1
<0.7
-

134
89
41
18
14
17

0.3
0.3
<0.2
<0.2

Track 4
5.1
1.7
0.4
0.7
0.5
0.2

0.4
0.2
0.2
-

Crystal 2
Track 5
1
2
3
4
5
6

11
5.3
2.7
1.1
0.8
0.5

453
820
503
101
50
35

2.4
0.5
1.2
0.4
0.2
<0.2

Track 6
1.8
3.6
0.7
0.5
0.4
0.5

13
9.1
5.9
4.7
3.1
1.3

855
845
521
365
206
89

2.9
1.1
0.9
0.2
1.0
0.3

Track 7
3.1
4.1
1.8
1.3
0.6
0.7

13
8.4
2.8
3.8
2.5
0.9

896
624
378
210
130
112

6.7
2.4
0.5
1.6
0.2
0.2

Track 8
1.2
0.4
0.4
<0.2
0.6
<0.2

3.0
3.0
2.8
0.7
0.5
1.1

The maximum PGE contents at the Natalkinskoe deposit are currently known to be associatedwith sulphide minerals. High levels of PGE in the sulphide gravity concentrate of this deposit ore
were established by sample decomposition using ﬂuoro oxidants [49] and amounted to (in ppm) Pt,
12.9–92.8; Pd, 0.62–2.97; Ru, 0.11–3.27; Rh, 0.95–1.54 and Ir, 0.1–0.3. The speculation about the possible
role of sulphide minerals as Pt concentrators at the Natalkinskoe gold deposit was previously made by
L.P. Plyusnina and co-authors [14]; however, microprobe analysis did not reveal high content of noble
metals in conventional preparations of arsenopyrite and pyrite [3]. Arsenopyrite concentrates of not
only Au, but also Pt and Pd, were later demonstrated at the Natalkinskoe deposit [11]. The new data
conﬁrmed and supplemented the earlier conclusion. The highest contents of PGE (in ppm) observed inarsenopyrite were Pt up to 128, Pd up to 20, Ru up to 86 and Rh up to 21, and they were less high inpyrite: Pt up to 29, Pd up to 15, Ru up to 58 and Rh up to 5.9. AAS-ADSSC found that the high content
of PGE largely was due to the existence of non-autonomous phases, causing their surface concentration.
The fact that the nature of high contents of PGE in arsenopyrite and pyrite is mainly superﬁcial was
conﬁrmed by the data obtained by SEM-EDX and (to some extent) LA-ICP-MS. If the data from theﬁrst two methods are comparable (taking into account the above considerations), the latter gives
diﬀerent results, which can be explained by the following reasons. First, there is some dispersion of
the material on the surface after the ﬁrst and each subsequent laser pass, which increases the error ofdetermination. Second, the analysis of surface areas of several crystals does not have the necessary
statistical reliability for determining the average contents, as is the case in the analytical technology
of AAS-ADSSC. It is quite possible to meet with an “empty” crystal or, conversely, a crystal with thesurface strongly enriched with impurity. This is all the more likely because the heterogeneity of the
distribution of PGE is well known. For example, in pyrite from the intercumulus sulphide phases ofthe main sulphide zone of Great Dyke (Zimbabwe), the ion-microprobe analysis revealed a strong
heterogeneity in the Pt distribution—from 0.4 to 244 ppm in diﬀerent grains of the same sample [50].
Thirdly, signiﬁcant errors can occur due to the lack of external calibration on such surface structures,for example, deposited on the crystal ﬁlms of sulphide material with a certain content of admixtureof the studied element. In the current state of the study of the surface composition of sulphides by
LA-ICP-MS, although the history of such studies seems to date back 12 years [51], it is possible tospeak only about the general tendency consisting of the enrichment of PGE of a surface of pyrite andarsenopyrite crystals; however, this trend is fully conﬁrmed by other methods.

91

Minerals 2020, 10, 318

The reason for this phenomenon involves the features of the crystal growth mechanism and the
dualism of the element distribution coeﬃcient in the mineral-hydrothermal solution system, which
is an order of magnitude higher for NAP compared to the crystal volume [38,39]. Non-autonomous
phases are extremely important, but they are still rarely considered components of synthetic materials
and minerals, the geochemical role of which lies in their ability to accumulate impurity elements inultra-high contents [52]. Such phases occur on the surface of a growing crystal due to the chemical
modiﬁcation and structural reconstruction of its surface layer, which is in local equilibrium with asupersaturated solution. NAP contain defects, chemical bonds and valence states of elements unusual
for crystal volume, for example AsV, AsIII, AsII and FeIII on pyrite [41], which provides additional
opportunities for the incorporation of minor and trace elements, especially incompatible, into theirstructures. Since the stability ﬁeld of arsenopyrite is adjacent to the pyrite ﬁeld, according to theprinciple of continuity of phase formation on mineral surfaces and the correspondence of chemical
forms of elements on them [53], the surface NAP on arsenopyrite can have a defective pyrite structure.
Since platarsite Pt(As,S)2 has the structure of pyrite Pa3, we should therefore expect an increase in thesolubility of Pt in the surﬁcial phase compared to the volume of the crystal.
Gold and PGE nanoparticles are increasingly, especially recently, found in ores of diﬀerent
genesis deposits [7,8,54–56]. Although no PGE mineral forms were identiﬁed at the Natalkinskoe
deposit, the need to continue research in this direction is obvious. This may be demonstrated by the
Degdekan deposit, another gold project in the northeast of Russia, which is in the territory of Pre
Kolyma. Its ores, classed as black shale formation, yielded the ﬁrst ﬁndings of a number of platinumgroup minerals of submicron dimensions: native osmium, routheniridosmine, osmirid, ruthenosmirid,
laurite, iridarsenite, ruthenium arsenide and arsenide of osmium and iridium [8]. The studies by [5,6]established the existence of diﬀerent mineral forms of platinum metals, including native platinum,
phases of Pt–Cu–Fe system, sperrylite, cooperite, and minerals of the Pd–Bi system in gold ores of the
Sukhoi Log deposit (Lena ore district, Irkutsk region, Russia). Their size mainly ranges within 1–10 μm.
Later, native Pt nanoparticles, with a dominant size of 0.5–20 nm, were found in the concentrates of
insoluble carbonaceous matter of shales containing ore mineralization [7]. It follows that there is ahigh probability of nanoscale mineral forms of PGE existing at the Natalkinskoe deposit, according
to the model of surface NAP evolution in sulphide minerals, suggesting their partial transformation
and aggregation with the formation of nano- and microinclusions of separate (autonomous) phases oftrace elements [39]. The possibility of ﬁnding nano-sized particles of PGE in sulphides is conﬁrmed byexperimental data. For example, nanometer-size PtS2 inclusions in synthetic Pt-containing pyrite weredetected with high-resolution transmission electron microscopy [57].
5. Conclusions
Arsenopyrites and pyrites of the Natalkinskoe gold deposit selected from vein, streaky-vein andveinlet-disseminated ores belonging to the most productive gold-rich hydrothermal mineral formations
concentrated not only Au, but also platinum group elements. The highest contents (in ppm) weredetected in monofractions of arsenopyrite: Pt up to 128, Pd up to 20, Ru up to 86 and Rh up to 21,
and less high in monofractions of pyrite: Pt up to 29, Pd up to 15, Ru up to 58 and Rh up to 5.9.
The original AAS-ADSSC analytical technology in arsenopyrites and pyrites from the deposit
established that there were two forms of uniformly distributed Pt, Pd, Ru and Rh—structural andsurface-bound chemically bound in the structure of these minerals and in non-autonomous phases
located on the surface of arsenopyrite and pyrite crystals. The surﬁcially bound form dominates andprobably exists in a very thin surface layer of the crystal (~100–500 nm). This phenomenon arises due tothe features of the crystal growth mechanism and the dualism of the element distribution coeﬃcient in
the mineral-hydrothermal solution system, which is an order of magnitude higher for NAP comparedto the crystal volume. It is assumed that for gold-bearing arsenopyrites and pyrites there is a common
mechanism of impurity absorption associated with the active role of the crystal surface and surface
defects. Structural forms of Pt, Pd, Ru and Rh, despite a lower content of the associated element

92Minerals 2020, 10, 318

compared to the surface, are also extremely important. They are a reliable indicator of element activity
in ore-forming ﬂuids forming gold deposits and a criterion of their potential ore content.
Restoration of the ore-forming ﬂuid composition using the content of the structural admixtureof Au and Pt in pyrite and experimentally established coeﬃcients of their distribution in the
pyrite–hydrothermal solution system showed that ore-forming ﬂuid that deposited pyrite locallycontained comparable concentrations of gold and platinum. Unlike gold, which mostly occurredin visible native form, the majority of PGEs are in so-called hidden or invisible forms. Previously,
they were not recognized due to the inability to establish the nature of the carrier without the use of
methods to determine the surﬁcial and structural forms of the element.
The data resulting from studying the surface of sulphide minerals from the Natalkinskoe deposit
using SEM-EDX and LA-ICP-MS methods conﬁrmed the surface presence on arsenopyrite and pyrite
crystals of Pt, Pd, Ru and Rh and such PGE as Os and Ir, as well as other ore trace elements including
Ag, Hg, Pb, Cu and U. The maximum contents of invisible forms of PGE, as in the case of gold,
are conﬁned to the surface of crystals. An early assumption that surface-bound forms, as the main
uniformly distributed form of elements in ore minerals, can be characteristic not only for Au, Pt and
Pd was conﬁrmed [11]. They are typical of other PGE (Ru, Rh, Os and Ir) and, in general, for ore
trace elements.
The contradiction of LA-ICP-MS data with other methods with respect to the distribution of Pt
and Pd in pyrite was revealed. This is assumed to be due to the presence of an unknown surface
form of Pd, inhomogeneously distributed over crystals (Pd nanoparticles or any of its compounds).
Although no proper mineral forms of PGE (Pt, Pd, Ru, Rh, Ir and Os) were established at the
Natalkinskoe gold deposit, the probability of nano- or micro-sized mineral forms of PGE at this deposit
is very high. This conclusion is based on the model of the evolution of surface NAP, assuming their
partial transformation and aggregation with the formation of nano- and microinclusions of separate(autonomous) phases of trace elements.
The results obtained through the study of PGEs in sulphide minerals are of interest boththeoretically and practically. Data on the nature of distribution, concentration level and forms of Pt,
Pd, Ru, Rh, Ir and Os in arsenopyrites and pyrites signiﬁcantly supplement existing ideas concerning
peculiarities in the composition of gold mineralization, and they serve as a criterion for assessing
the potential ore content of ﬂuids forming gold deposits. The association of maximum contents ofinvisible forms of PGEs with the surface of sulphide mineral crystals can be used to develop methods of
PGE recovery, primarily Pt, without destruction of the mineral structure, through transformation anddissolution of the surface layer only. This is an important point that must be taken into account in theconcentration of ores, which includes sulphides, and in our case arsenopyrite and pyrite. The presenceof PGE in ores and the possibility of their extraction signiﬁcantly enhances the value of the gold ore
mined, thereby increasing the prospects of the deposit under study.
Author Contributions: Conceptualization, R.G.K.; methodology, R.G.K. and V.L.T.; validation, R.G.K.;
formal analysis, V.L.T., A.S.M., N.V.B. and N.V.S.; investigation, R.G.K. and V.L.T.; resources, R.G.K. and
A.S.M.; writing—original draft preparation, R.G.K. and V.L.T.; writing—review and editing, R.G.K., V.L.T., A.S.M.
and N.V.B.; visualization, R.G.K., A.S.M. and N.V.S. All authors have read and agreed to the published version of
the manuscript.
Funding: The work was carried out within the framework of the execution of the state task for the projects
IX.125.3.4. (0350-2019-0003) and IX.130.3.1. (0350-2019-0010) with the ﬁnancial support of the RFBR (projects Nos.
20-05-00142 and 18-05-00077) and using scientiﬁc equipment of the shared research centre “Isotopic-geochemical
investigations” IGC SB RAS (Irkutsk, Russia) and the shared research centre “Ultramicroanalysis” LIN SB RAS(Irkutsk, Russia).
Acknowledgments: The authors thank T.M. Pastushkova, I.Yu. Voronova, V.N. Vlasova, G.A. Shcherbakova, S.V.
Lipko and K.Yu. Arsent’ev for their assistance in analytical studies. We would like to express our special gratitude
to all geological services of the JSC “RiM” (Magadan, Russia), its chief geologist R.N. Ovsov and leading geologist
E.M. Nikitenko, for comprehensive assistance in conducting ﬁeldwork in the period from 2008 to 2016.
Conﬂicts of Interest: The authors declare no conﬂict of interest.

93

Minerals 2020, 10, 318

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Direct imaging of gold nanoparticules in a Carlin-type deposit. Am. Mineral. 2004, 89, 1359–1366. [CrossRef]
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chromite deposits. Ore Geol. Rev. 2017, 81, 1236–1248. [CrossRef]
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Vikentyev, I.V.; Tagirov, B.R. The state of platinum in pyrite studied by X-ray absorption spectroscopy ofsynthetic crystals. Econ. Geol. 2019, 114, 1649–1663. [CrossRef]© 2020 by the authors. Licensee MDPI, Basel, Switzerland. This article is an open accessarticle distributed under the terms and conditions of the Creative Commons Attribution(CC BY) license (http://creativecommons.org/licenses/by/4.0/).

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minerals
ArticleZoned Laurite from the Merensky Reef, Bushveld
Complex, South Africa: “Hydrothermal” in Origin?
Federica Zaccarini * and Giorgio Garuti
Department of Applied Geological Sciences and Geophysics, University of Leoben, A-8700 Leoben, Austria;
giorgio.garuti1945@gmail.com* Correspondence: federica.zaccarini@unileoben.ac.at; Tel.: +43-(0)3842-402-6218
Received: 21 March 2020; Accepted: 19 April 2020; Published: 21 April 2020

Abstract: Laurite, ideally (Ru,Os)S2 , is a common accessory mineral in podiform and stratiform
chromitites and, to a lesser extent, it also occurs in placer deposits and is associated with Ni-Cu
magmatic sulﬁdes. In this paper, we report on the occurrence of zoned laurite found in the Merensky
Reef of the Bushveld layered intrusion, South Africa. The zoned laurite forms relatively large crystals
of up to more than 100 μm, and occurs in contact between serpentine and sulﬁdes, such as pyrrhotite,chalcopyrite, and pentlandite, that contain small phases containing Pb and Cl. Some zoned crystalsof laurite show a slight enrichment in Os in the rim, as typical of laurite that crystallized at magmaticstage, under decreasing temperature and increasing sulfur fugacity, in a thermal range of about
1300–1000 ◦ C. However, most of the laurite from the Merensky Reef are characterized by an unusual
zoning that involves local enrichment of As, Pt, Ir, and Fe. Comparison in terms of Ru-Os-Ir ofthe Merensky Reef zoned laurite with those found in the layered chromitites of the Bushveld andpodiform chromitites reveals that they are enriched in Ir. The Merensky Reef zoned laurite also
contain high amount of As (up to 9.72 wt%), Pt (up to 9.72 wt%) and Fe (up to 14.19 wt%). On thebasis of its textural position, composition, and zoning, we can suggest that the zoned laurite of theMerensky Reef is “hydrothermal” in origin, having crystallized in the presence of a Cl- and As-richhydrous solution, at temperatures much lower than those typical of the precipitation of magmatic
laurite. Although, it remains to be seen whether the “hydrothermal” laurite precipitated directly fromthe hydrothermal ﬂuid, or it represents the alteration product of a pre-existing laurite reacting withthe hydrothermal solution.
Keywords: laurite; sulﬁdes; ﬂuids; platinum group elements (PGE); platinum group minerals (PGM);
Merensky Reef; Bushveld Complex; South Africa

1. Introduction
Minerals of ruthenium are very rare and only ﬁve of them, namely anduoite (Ru,Os)As2 , laurite(Ru,Os)S2 , ruarsite RuAsS, ruthenarsenite (Ru,Ni)As, and ruthenium (Ru,Ir,Os), have been approvedby the International Mineralogical Association (IMA). They occur as accessory minerals associatedwith maﬁc–ultramaﬁc rocks, especially with chromitite, and as nuggets in placer deposits. Amongthe minerals of ruthenium, laurite is the most common. It was discovered in 1866 in a placer from
Laut, Banjar, South Kalimantan Province, Borneo, Indonesia [1]. Laurite is a common constituentof the suite of platinum group minerals (PGM) inclusions (usually less than 20 μm) in podiform
and stratiform chromitites [2–4]. Less frequently, laurite has been reported from placers and Ni-Cu
magmatic sulﬁde deposits [5–7]. Laurite forms a complete solid solution with erlichmanite (OsS2 ) [5],and their typical mode of occurrence, i.e., included in chromite grains, indicate that they crystallized athigh temperatures, in a thermal range of about 1300–1000 ◦ C prior to, or coeval with, the precipitationof the host chromitite [2–4]. The reciprocal stability of laurite and erlichmanite is strongly controlled

Minerals 2020, 10, 373; doi:10.3390/min10040373

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by sulfur fugacity and temperature. In particular, laurite precipitates at a higher temperature andlower sulfur fugacity, compared to erlichmanite. This order of crystallization can be observed in thezoning of the small crystals of laurite and erlichmanite enclosed in fresh chromite grains that, typically,show an Os-poor core, grading into a high-Os rim [2–4]. This magmatic zoning can be obliterated
by low temperature processes such as serpentinization and weathering, as documented in laurite
associated with podiform chromitites [8–10]. During alteration processes at low temperature, laurite
and erlichmanite lose their original S and release part of Os and Ir to form secondary Ru-Os-Ir alloys, inwhich the lost S may be replaced by Fe-oxide [8–10]. The occurrence of laurite in the Bushveld Complex
of South Africa has been documented by several authors [2,7,11–17]. Most commonly, the mineraloccurs as small polygonal grains enclosed in chromite grains of the Critical Zone chromitite layers andhas only occasionally been found as part of the sulﬁde ore of the Merensky Reef. In this contribution,
we have investigated in detail the mineral chemistry of the laurite associated with the sulﬁde-rich zoneof the Merensky Reef. The grains are characterized by an unusual zoning and composition comparedwith laurite inclusions in the Bushveld chromitites, suggesting that the mineral was generated underdiﬀerent thermodynamic conditions in the two cases.
2. Sample Provenance and Petrographic Description
The Bushveld layered intrusion is located in the central part of the Transvaal province, north ofPretoria, South Africa (Figure 1A), and it is divided into Eastern, Western, and Northern limbs (Figure 1B).
The Bushveld intrusion is well known among economic geologists because it contains the world’s
largest deposits of platinum group elements (PGE), namely: the UG-2 chromitite and the Merensky
Reef [18]. The noun Merensky Reef refers to a sulfide-bearing pegmatoidal feldspathic pyroxenite
enriched in PGE, marked at the bottom and top, by two centimetric layers of chromitite. The Reef can betraced for a total strike of about 280 km, marking the limit between the Critical and Main Zone [18,19].
The investigated samples were collected by one of the authors (G.G.) in the Rustenburg undergroundmine, during the third International Platinum Symposium held in Pretoria from 6 to 10 July 1981 [20,21].
The Rustenburg mine is located in the Western limb of the Bushveld Complex, about 100 km westof Pretoria (Figure 1B,C). Here, four different zones of Bushveld (undifferentiated in Figure 1C) are
intruded by the Pilanesberg Alkaline Complex [22]. Four square polished blocks, about 2.5 × 2.5 cm(Figure 2), were prepared for petrographic and mineralogical investigation. The blocks consist of a thinlayer of chromitite, about 0.2 cm thick, in contact with pegmatoidal feldspathic pyroxenite and large
blebs of sulfide. In agreement with observations made by several authors [21,23,24], the sulfide-richzone contains accessory actinolite, micas, talc, chlorite, and a serpentine subgroup mineral.
3. Methods
The polished blocks were previously studied by reflected-light microscope. Quantitative chemical
analyses of laurite were performed with a JEOL JXA-8200 electron microprobe (JEOL, Tokyo, Japan),installed in the E. F. Stumpfl laboratory, Leoben University, Leoben, Austria, operated in WDS(wavelength dispersive spectrometry) mode. Major and minor elements were determined at 20 kV
accelerating voltage and 10 nA beam current, with 20 s as counting time for the peak and 10 s forthe backgrounds. The beam diameter was about 1 μm in size. The Kα lines were used for S, As, Fe
and Ni; Lα for Ir, Ru, Rh, Pd, and Pt, and Mα for Os. The reference materials were pure metals forthe six PGE (Ru, Rh, Pd, Os, Ir and Pt), synthetic NiS, natural pyrite and niccolite for Ni, Fe, S and
As. The following diffracting crystals were selected: PETJ for S; PETH for Ru, Os, Pd and Rh; LIFH
for Fe, Ni, Ir and Pt; and TAP for As. Automatic correction was performed for the Ru-Rh and Rh-Pd
interferences. The detection limits were calculated by the software and are: Os (0.07 wt%), Ir (0.06 wt%),
Ru, Pd, and Pt (0.04 wt%), Rh (0.03 wt%), Fe, Ni, As and S (0.02 wt%). The grains smaller than 10 μm
were analyzed by EDS. The same instrument was used to obtain back-scattered electron images (BSE)and X-ray elemental distribution maps.

98Minerals 2020, 10, 373

Figure 1. The Bushveld Complex, South Africa (A) and locations of the Rustenburg mine and theMerensky Reef (B,C) in the Western limb (modiﬁed after [15,18]).

Figure 2. Example of the studied polished blocks from the Merensky Reef, showing the sulﬁde blebs(creamy–yellow) and the cumulitic chromitite (small polygonal dark grey grains) in the pegmatoidal
feldspathic pyroxenite.

4. Laurite: Morphology, Texture, and Composition
The investigated samples contain laurite in two diﬀerent textural positions, either included in
fresh chromite of the thin chromitite layer (Figure 3A), or at the contact between sulﬁde patches(pyrrhotite, chalcopyrite, pentlandite) and serpentine (Figure 3B–D).

99

Minerals 2020, 10, 373

Figure 3. Digital image in reﬂected plane polarized light showing laurite from the Merensky Reef. (A)Laurite enclosed in fresh chromite. (B,C) Zoned laurite and (D) enlargement of (C). Abbreviations: Lrt= laurite, Plg = plagioclase, Chr = chromite, Srp = serpentine, Pn = pentlandite, Po = pyrrhotite, Chp =chalcopyrite. Scale bar = 20 μm.

Quantitative analyses of laurite enclosed in chromite and associated with sulﬁdes are listed inTables 1 and 2, respectively. Laurite included in chromite forms tiny crystals, usually not exceeding
10 μm in size, characterized by euhedral to subeuhedral morphology and homogenous composition.
Laurite associated with sulﬁdes and serpentine is bigger, up to more than 100 μm, and may occur as
single crystals or clusters of grains (Figure 3B–D), characterized by subeuhedral to anhedral shape.
The BSE images of large laurite display remarkable zoning emphasized by marked contrast in theelectronic reﬂectivity (Figure 4A–D). Laurite in the sulﬁde assemblage is accompanied by a complex
association of precious minerals comprising: cooperite (PtS), moncheite (PtTe2 ), platarsite (PtAsS),
rustenburgite (Pt3 Sn), Pt-Fe alloy, undetermined Pt-Te-Bi and Pd-Te-Bi compounds, Au-Ag alloy, andthe recently discovered PGM bowlesite PtSnS [21].
Table 1. Selected wavelength dispersive spectrometry (WDS) electron microprobe analyses of Merensky
Reef laurite enclosed in chromite.
Sample

As

S

Ru

Os

Ir

Rh

Pt

Pd

Ni

Fe

Total

1.24
1.06

0.00
0.00

1.79
1.68

0.03
0.08

1.00
0.90

101.18100.48

Rh

Pt

Pd

Ni

Fe

0.69
0.59

0.00
0.00

0.96
0.90

0.03
0.08

1.02
0.91

wt%
mr18a
mr8a

1.69
1.85

36.81
36.93

50.36
49.99

5.42
4.97

2.85
3.03

Sample

As

S

Ru

Os

Ir
at%

mr18a
mr8a

1.28
1.40

65.25
65.58

28.32
28.16

1.62
1.49

0.84
0.90

100

Minerals 2020, 10, 373

Table 2. WDS electron microprobe analyses of Merensky Reef zoned laurite.
Sample

As

S

Ru

Os

Ir

Rh

Pt

Pd

Ni

Fe

Total

1.09
0.90
1.38
1.32
1.35
1.37
1.19
1.63
1.05
1.16
1.53
1.75
2.02
1.98
1.77
1.08
1.12
1.18
1.38
1.36
1.11
1.23
1.32
0.82
3.03
3.50
2.69
2.75
2.80
2.68
2.51
2.54
1.18
1.22
1.09
1.10
1.14
1.31
1.09
1.56
1.21
1.24
1.06

0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.10
0.00
0.00
7.20
1.78
1.04
0.85
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.57
2.53
8.69
3.79
5.61
7.33
8.20
2.95
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.00

1.32
1.18
1.35
1.55
1.47
1.41
1.47
1.58
1.42
1.45
1.63
1.05
1.58
1.55
1.48
1.39
1.37
1.75
1.42
1.49
1.38
1.44
1.48
1.55
1.49
1.61
1.37
1.58
1.50
1.36
1.41
1.62
1.40
1.35
1.45
1.43
1.22
1.24
1.31
1.45
1.10
1.79
1.68

0.26
0.06
0.42
0.81
0.90
0.65
0.49
0.39
0.53
0.31
0.57
0.80
0.63
0.54
0.76
0.42
0.34
0.31
0.35
0.27
0.41
0.22
0.39
0.82
0.23
0.30
0.36
0.35
0.31
0.35
0.31
0.27
0.56
0.56
0.71
0.52
0.12
0.13
0.73
0.67
1.39
0.03
0.08

6.54
5.76
5.87
4.23
4.73
5.22
5.24
5.44
5.28
6.17
4.57
3.05
4.35
4.75
5.57
5.26
5.74
6.26
6.38
5.78
5.51
5.84
6.40
4.61
14.19
10.13
5.66
7.79
7.78
5.90
5.97
8.77
4.03
3.82
3.70
3.94
5.98
8.68
3.75
4.13
4.95
1.00
0.90

99.13
100.81
100.27101.15100.7399.94
100.21100.73100.13101.66100.23
99.57
101.20100.91101.02101.14100.63101.24100.8099.84
100.97100.46100.63101.20100.58100.2799.38
100.8399.56
100.90100.44101.17101.19101.3799.51
101.21100.59100.92101.38100.41
100.16101.18100.48

wt%
MR2a1
MR2a2
MR2a3
MR2a4
MR2a5
MR2a6
MR2a7
MR2a8
MR2a9
MR2a10
MR2a11
MR2a12MR2a13
MR2a14
MR2a15
MR2a16
MR2a17
MR2a18MR2a19MR2a20MR2a21MR2a22MR2a23
MR2a24
MR4a1
MR4a2
MR4a3
MR4a4
MR4a5
MR4a6
MR4a7
MR4a8
MR1a1
MR1a2
MR1a3
MR1a4
MR1a5
MR1a6
MR1a7
MR1a8
MR1a9
MR1a10
MR1a11

1.08
1.28
1.63
1.51
1.38
1.74
1.44
1.42
1.50
1.54
2.22
9.72
3.66
2.90
2.74
1.39
1.41
1.49
1.40
1.42
1.42
1.49
1.50
1.54
2.06
3.90
8.19
4.32
5.78
6.88
7.94
4.26
1.38
1.36
1.44
1.43
1.45
1.46
1.45
2.03
1.51
1.69
1.85

37.50
36.97
37.63
37.58
37.68
37.32
38.69
39.13
38.46
38.98
38.17
28.45
36.78
37.25
36.97
39.25
39.21
39.68
39.70
38.14
39.38
39.29
39.63
37.92
41.46
38.45
32.52
37.85
34.99
34.58
32.54
37.60
37.10
37.54
36.34
37.43
37.60
38.23
36.94
36.96
37.75
36.81
36.93

39.32
37.20
43.23
45.24
43.61
42.99
44.54
44.65
43.86
43.03
44.63
28.24
42.64
43.43
41.61
44.22
44.41
44.34
43.20
41.29
44.15
45.22
43.90
43.50
33.93
35.37
33.93
37.63
34.15
36.72
33.50
36.69
42.11
42.06
42.08
41.91
39.58
37.79
41.04
41.34
36.32
50.36
49.99

4.79
7.97
2.53
3.15
3.59
2.78
1.88
2.00
1.86
2.77
1.57
5.90
1.90
1.64
2.55
2.38
1.92
1.70
1.73
3.31
2.30
1.47
1.38
3.35
0.94
1.51
3.15
2.52
3.55
2.47
3.92
3.21
4.44
4.85
4.46
4.47
6.03
5.25
7.93
2.77
7.72
5.42
4.97

7.23
9.49
6.22
5.75
6.02
6.46
5.27
4.50
6.06
6.25
5.35
13.43
5.85
5.83
6.73
5.75
5.11
4.52
5.22
6.79
5.31
4.26
4.63
7.10
2.68
2.98
2.83
2.25
3.08
2.64
4.15
3.26
9.00
8.61
8.24
8.97
7.48
6.84
7.14
9.50
8.20
2.85
3.03

101

Minerals 2020, 10, 373

Table 2. Cont.
Sample

As

S

Ru

Os

Ir

Rh

Pt

Pd

Ni

Fe

0.59
0.50
0.74
0.71
0.73
0.74
0.63
0.86
0.56
0.61
0.82
1.11
1.10
1.07
0.96
0.57
0.59
0.61
0.72
0.73
0.58
0.64
0.68
0.44
1.49
1.82
1.55
1.45
1.55
1.49
1.45
1.34
0.65
0.66
0.61
0.60
0.62
0.70
0.60
0.86
0.66
0.69
0.59

0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.03
0.00
0.00
2.41
0.51
0.30
0.24
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.15
0.69
2.65
1.06
1.64
2.15
2.49
0.82
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.00

0.70
0.63
0.70
0.81
0.76
0.74
0.75
0.80
0.73
0.73
0.84
0.64
0.83
0.81
0.77
0.71
0.69
0.88
0.72
0.78
0.70
0.73
0.74
0.80
0.71
0.81
0.77
0.81
0.80
0.73
0.79
0.82
0.74
0.72
0.78
0.76
0.64
0.63
0.70
0.77
0.58
0.96
0.90

0.25
0.05
0.39
0.77
0.85
0.62
0.45
0.36
0.50
0.28
0.53
0.89
0.60
0.51
0.72
0.39
0.31
0.28
0.32
0.25
0.38
0.20
0.35
0.78
0.20
0.28
0.37
0.32
0.30
0.34
0.31
0.25
0.54
0.54
0.70
0.50
0.11
0.12
0.71
0.65
1.33
0.03
0.08

6.58
5.88
5.82
4.20
4.70
5.21
5.13
5.26
5.19
5.98
4.50
3.57
4.35
4.72
5.55
5.10
5.56
5.98
6.11
5.73
5.32
5.63
6.13
4.56
12.84
9.68
6.02
7.58
7.94
6.04
6.34
8.52
4.08
3.85
3.81
3.98
6.00
8.49
3.81
4.19
5.00
1.02
0.91

at%
MR2a1
MR2a2
MR2a3
MR2a4
MR2a5
MR2a6
MR2a7
MR2a8
MR2a9
MR2a10
MR2a11
MR2a12MR2a13
MR2a14
MR2a15
MR2a16
MR2a17
MR2a18MR2a19MR2a20MR2a21MR2a22MR2a23
MR2a24
MR4a1
MR4a2
MR4a3
MR4a4
MR4a5
MR4a6
MR4a7
MR4a8
MR1a1
MR1a2
MR1a3
MR1a4
MR1a5
MR1a6
MR1a7
MR1a8
MR1a9
MR1a10
MR1a11

0.81
0.97
1.20
1.12
1.02
1.29
1.05
1.02
1.10
1.11
1.63
8.48
2.73
2.15
2.03
1.01
1.02
1.06
1.00
1.05
1.03
1.07
1.07
1.14
1.39
2.78
6.50
3.13
4.40
5.25
6.29
3.09
1.04
1.02
1.11
1.08
1.08
1.07
1.10
1.53
1.14
1.28
1.40

65.69
65.76
64.94
64.99
65.21
64.96
65.89
65.97
65.81
65.72
65.43
58.03
64.07
64.45
64.14
66.26
66.10
66.04
66.30
65.91
66.27
66.01
66.11
65.45
65.32
64.02
60.32
64.09
62.15
61.69
60.18
63.62
65.42
65.85
65.22
65.77
65.67
65.12
65.50
65.24
66.34
65.25
65.58

21.85
20.99
23.67
24.82
23.94
23.74
24.06
23.89
23.81
23.02
24.27
18.27
23.56
23.84
22.90
23.68
23.75
23.41
22.89
22.63
23.58
24.10
23.23
23.82
16.96
18.68
19.97
20.21
19.24
20.78
19.66
19.69
23.56
23.41
23.96
23.36
21.93
20.43
23.09
23.15
20.25
28.32
28.16

1.41
2.39
0.74
0.92
1.05
0.82
0.54
0.57
0.54
0.79
0.45
2.03
0.56
0.48
0.75
0.68
0.55
0.48
0.49
0.96
0.65
0.42
0.39
0.97
0.25
0.42
0.98
0.72
1.06
0.74
1.22
0.91
1.32
1.43
1.35
1.32
1.77
1.51
2.37
0.82
2.29
1.62
1.49

2.11
2.82
1.79
1.66
1.74
1.88
1.50
1.26
1.73
1.76
1.53
4.57
1.70
1.68
1.95
1.62
1.44
1.25
1.45
1.96
1.49
1.19
1.29
2.04
0.71
0.83
0.88
0.63
0.91
0.78
1.28
0.92
2.65
2.52
2.47
2.63
2.18
1.94
2.11
2.80
2.41
0.84
0.90

102

Minerals 2020, 10, 373

Figure 4. Back-scattered electron (BSE) images of zoned laurite. (A) See reﬂected-light image of
Figure 3B for the mineralogical assemblage of the grain. (B) Laurite in contact with pyrrhotite andserpentine, (C) laurite in contact with chalcopyrite and serpentine and (D) laurite grains in contact
with pyrrhotite and serpentine Abbreviations as in Figure 3. Scale bar = 20 μm.

As previously reported by [14], abundant Pb-Cl minerals, less than 10 μm in size, were also
observed enclosed in the sulﬁdes (Figure 5A), and qualitatively identiﬁed by EDS (Figure 5B). The EDS
overlap between Pb and S was checked by a WDS semi-quantitative analysis that gave a composition(wt%) of 77.8 for Pb and 18.9 for Cl, very similar to the mineral analyzed by [14].

Figure 5. (A) BSE image of a Pb-Cl mineral (Pb-Cl) enclosed in pyrrhotite (Po), scale bar = 10 μm, and(B) its EDS spectrum (see the text for the Pb-S overlap).

Electron microprobe analyses of the zoned laurite (Table 2) and elemental distribution maps
(Figures 6–8) showed unusual enrichments in As, Ir, Os, Pt, and Fe. Distribution of Rh, Pd, and Ni was
not visible in the X-ray maps because of the low concentrations, while Cu (not analyzed, but visible in
Figure 7) was due to a Cu-phase ﬁlling ﬁssures in laurite. Substitution of As for S occurs systematicallyfrom a homogeneous background of about 1.00–1.50 wt% (Figure 6) up to patchy enrichment of
3.66–9.7 wt% (Figures 7 and 8). The enrichments of Os and Ir are closely related and may occur either
at the rim of grains as described by [14] (Figure 6), or as irregular patches (Figures 7 and 8). The Pt
appears to be particularly concentrated, up to 8.69 wt%, in the As-rich zones (Figures 7 and 8).

103Minerals 2020, 10, 373

Figure 6. BSE image and X-ray element-distribution maps of S, As, Ru, Ir, Os, Pt, Cu, and Fe in zoned
laurite from the Merensky Reef. See Figures 3B and 4A for the paragenetic assemblage. Scale bar =20 μm.

Figure 7. BSE image and X-ray element-distribution maps of S, As, Ru, Ir, Os, Pt, Cu, and Fe, showing
the zoning of the laurite from the Merensky Reef, see reﬂected-light image of Figure 4D for themineralogical assemblage. Scale bar is 20 μm.

104Minerals 2020, 10, 373

Figure 8. BSE image and X-ray element-distribution maps of S, As, Ru, Ir, Os, Pt, Cu, and Fe, showing
the zoning of the laurite from the Merensky Reef, see reﬂected-light image of Figure 4C for themineralogical assemblage. Scale bar is 20 μm.

The Ru-Os-Ir (wt%) ternary diagram (Figure 9) shows that the zoned laurite of the Merensky Reef
are significantly enriched in Ir, compared with laurite enclosed in the chromitite of the same Reef, andother chromitite layers of the Bushveld. They also do not display the Ru-Os negative correlation inferredby the Ru-Os substitution trend due to the laurite-erlichmanite solid solution trend (Figure 8). Basedon more than 1000 published analyses, and unpublished data of the authors, laurite associated with
ophiolitic, stratiform, and Alaskan-type magmatic chromitites exhibit a pronounced negative correlation
between Ru and Os (R = −0.97). In contrast, the correlation matrix calculated from our electron
microprobe analyses (Table 3) indicates that zoned laurite of the Merensky Reef are characterized by theabsence of Ru-Os correlation (R = −0.07).
Table 3. Element correlation in the zoned laurite for the Merensky Reef.
at%

As

S

Ru

Os

Ir

Rh

Pt

Pz

Ni

Fe

As
S
Ru
Os
Ir
Rh
Pt
Pd
Ni
Fe

1.00
−0.98
−0.56
0.10
0.04
0.72
0.97
0.03
0.07
0.11

1.00
0.53
−0.14
−0.09
−0.70
−0.95
−0.05
−0.09
−0.10

1.00
−0.07
−0.04
−0.68
−0.58
0.45
−0.02
−0.78

1.00
0.56
−0.25
0.06
−0.37
0.20
−0.35

1.00
−0.42
−0.10
−0.50
0.50
−0.37

1.00
0.75
0.20
−0.12
0.56

1.00
0.03
−0.01
0.18

1.00
−0.21
−0.27

1.00
−0.29

1.00

Note: the relevant correlations are highlighted in bold.

105

Minerals 2020, 10, 373

Figure 9. (A) Ru-Os-Ir ternary diagram (wt%) for magmatic laurite enclosed in the Lower-, Middle-,
and Upper-group chromitite layers: from the Bushveld Complex (yellow ﬁeld) and podiform chromitite(blue ﬁeld). Compositional ﬁelds after [11,12] and unpublished data of the authors. (B) Merensky Reef
zoned laurite (open square) and laurite enclosed in the Merensky Reef chromite (red square).

In addition, the high concentrations of As (up to 9.72 wt%), Pt (up to 9.72 wt%), and Fe (up to
14.19%) (Table 2), distinguish the zoned laurite of the Merensky Reef from the laurite inclusions in
diﬀerent types of chromitite (ophiolitic, stratiform, Alaskan-type).
The correlation matrix (Table 3) and distribution X-ray maps (Figure 7) clearly support a positive
correlation between Pt-As (R = +0.97), and a negative correlation of both elements with S (R = −0.95
and −0.98, respectively). The possible existence of submicroscopic inclusions of sperrylite (PtAs2 ) or
platarsite (PtAsS) in laurite was carefully checked, and discarded.
The fact that the sum of S+As atoms is consistently close to stoichiometry (S + As = 2.00) supportsthat Pt and As are parts of the laurite structure. Notably also the high concentrations of Fe are not due
to exotic inclusions, but Fe appears to be the major substitute for Ru, showing a negative correlation of
R = −0.78 (Table 3), that is a clear discrepancy with common magmatic laurite in general.
5. Origin of the Zoned Laurite in the Merensky Reef
Several theories for the origin of the Merensky Reef have been proposed, and they have been
recently summarized by [25]. The proposed genetic models include: (i) gravitational settling of crystals
that precipitated in the magma chamber during the orthomagmatic stage; (ii) hydrodynamic sorting ofa mobilized cumulate slurry in a large magma chamber, under slow cooling; (iii) crystallization at the
crystal mush-magma interface caused by a replenishment event; (iv) interaction of a hydrous meltwith a partially molten cumulate assemblage; (v) intrusion of magmas as sills into undiﬀerentiatednorite; and (vi) intrusion of magma into a pre-existing cumulate pile ([25] and references therein).
To explain the formation of the coarse-grained feldspathic orthopyroxenite enriched in PGE,
and associated chromitite in the Merensky Reef, several authors have invoked the reaction between
a late-stage hydrous melt with an unconsolidated cumulate assemblage [23–27]. On the basis ofnatural observations supported by experimental results, formation of tiny Os-Ir-Ru PGM inclusions
in chromitite can be modeled by a sequence of crystallization events controlled by sulfur fugacity
and temperature (T) [2–4]. The sulfur fugacity is expected to increase with decreasing T in magmatic
systems between about 1300–1000 ◦ C, and a consistent order of crystallization can be observed. Alloys
in the system Os-Ir-Ru are the ﬁrst to precipitate, followed by laurite, usually characterized by acore-to-rim increment of Os content, and ﬁnally, erlichmanite. Incorporation of IrS2 molecules in thelaurite structure is generally low, controlled by Ir activity in the system. However, the systematic
substitution of Os for Ru can be remarkable, and the composition can enter the ﬁeld of erlichmanite if
sulfur fugacity increases suﬃciently during magmatic crystallization.
106

Minerals 2020, 10, 373

At a ﬁrst instance, the Os enrichment observed at the rim of some zoned laurite of the Merensky
Reef may correspond to such a magmatic crystallization trend. However, other characteristics of theMerensky Reef zoned laurite, such as the unusual zoning that involves local enrichment of As, Pt, Ir,
and Fe (Figures 7 and 8), and the absence of Ru-Os substitution, are in apparent contrast with thisconclusion. The observed zoning also does not reﬂect a ﬂuctuation of the sulfur fugacity, suggested to
explain the oscillatory zoning of Ru and Os, described in the laurite from the Penikat Layered Complexof Finland [28]. The presence of abundant minerals containing Pb and Cl and occurring enclosed in thesulﬁdes associated with the zoned laurite indicates the presence of Cl in the system [14]. Accordingto [14], the Cl-rich phase precipitated from a late-stage solution or formed as a result of replacement ofa precursor galena by an aqueous hydrochloric solution in the ﬁnal stage of hydrothermal alteration, at
low-temperature. Theoretical and experimental work coupled with natural observations suggest thatboth Cl and As may be important for the transport and mobilization of the PGE during metasomaticand hydrothermal events [23,28–32].
The textural position and the coarse grain size of laurite crystals, as well as their paragenesis
including hydrous silicates, suggest crystallization at a late stage from a volatile-rich melt enriched in
As and Cl, after coalescence of an immiscible sulﬁde liquid. The close stabilization of serpentine points
to relatively low temperatures for the precipitation of zoned laurite, certainly much lower than thoserequired for the crystallization of tiny laurite included in the chromite seams of the Merensky Reef.
Although it is not possible to provide a precise temperature for the genesis of the zoned laurite in the
Merensky Reef, we suggest they were in the range of 400–200 ◦ C, similar to temperatures calculated for
PGM deposition in the hydrothermal Waterberg platinum deposit of Mookgophong, South Africa [31].
Therefore, we can suggest that the zoned laurite of the Merensky Reef is “hydrothermal” in origin,
having crystallized in the presence of a Cl- and As-rich hydrous solution, at temperatures much lower
than those typical of the precipitation of magmatic laurite. Although, it remains to be seen whether the“hydrothermal” laurite precipitated directly from the hydrothermal ﬂuid, or it represents the alteration
product of a pre-existing laurite reacting with the hydrothermal solution.
Author Contributions: F.Z. and G.G. wrote the manuscript and provided support in the data interpretation. G.G.
collected the studied sample, and F.Z. performed the chemical analyses. All authors have read and agreed to the
published version of the manuscript.
Funding: The authors are grateful to the University Centrum for Applied Geosciences (UCAG) for the access to
the E. F. Stumpﬂ electron microprobe laboratory.
Acknowledgments: Many thanks are due to the editorial staﬀ of Minerals, to the guest editor Maria EconomouEliopoulos, and two referees for their useful comments. We are honored to dedicate this manuscript to the memoryof our friend and colleague, Demetrios G. Eliopoulos.
Conﬂicts of Interest: The authors declare no conﬂicts of interest.

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2019-079. CNMNC Newsletter No. 52. Mineral. Mag. 2019, 83. [CrossRef]
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152, 1–11. [CrossRef]
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Nicholson, D.M.; Mathez, E.A. Petrogenesis of the Merensky Reef in the Rustenburg section of the Bushveld
Complex. Contrib. Mineral. Petrol. 1991, 107, 293–309. [CrossRef]
Hunt, E.J.; Latypov, R.; Horváth, P. The Merensky Cyclic Unit, Bushveld Complex, South Africa: Reality or
Myth? Minerals 2018, 8, 144. [CrossRef]
Mathez, E.A. Magmatic metasomatism and formation of the Merensky Reef, Bushveld Complex. Contrib.
Mineral. Petrol. 1995, 119, 277–286. [CrossRef]
Boudreau, A.E. Modeling the Merensky Reef, Busvheld Complex, Republic of South Africa. Contrib. Mineral.
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Barkov, A.; Fleet, M.E.; Martin, R.F.; Alapieti, T.T. Zoned sulﬁdes and sulfarsenides of the platinum-groupelements from the Penikat layered complex, Finland. Can. Mineral. 2004, 42, 515–537. [CrossRef]
Boudreau, A.E. Chlorine as an exploration guide for the platinum-group elements in layered intrusions.
J. Geochem. Explor. 1993, 48, 21–37. [CrossRef]
Kislov, E.V.; Konnikov, E.G.; Orsoev, D.; Pushkarev, E.; Voronina, L.K. Chlorine in the Genesis of the
Low-Sulﬁde PGE Mineralization in the Ooko-Dovyrenskii Layered Massif. Geokhimiya 1997, 5, 521–528.
Oberthür, T.; Melcher, F.; Fusswinkel, T.; van den Kerkhof, A.M.; Sosa, G.M. The hydrothermal Waterberg
platinum deposit, Mookgophong (Naboomspruit), South Africa. Part 1: Geochemistry and ore mineralogy.
Mineral. Mag. 2018, 82, 725–749. [CrossRef]
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Econ. Geol. 2015, 110, 505–530. [CrossRef]© 2020 by the authors. Licensee MDPI, Basel, Switzerland. This article is an open accessarticle distributed under the terms and conditions of the Creative Commons Attribution(CC BY) license (http://creativecommons.org/licenses/by/4.0/).

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minerals
ArticleGeochemistry of Rare Earth Elements in Bedrock
and Till, Applied in the Context of Mineral Potential
in Sweden
Martiya Sadeghi *, Nikolaos Arvanitidis and Anna Ladenberger
Department of Mineral Resources, Geological Survey of Sweden, Box 670, 751 28 Uppsala, Sweden;
Nikolaos.arvanitidis@sgu.se (N.A.); Anna.Ladenberger@sgu.se (A.L.)* Correspondence: martiya.sadeghi@sgu.se; Tel.: +46-1817-9232
Received: 15 March 2020; Accepted: 15 April 2020; Published: 18 April 2020

Abstract: The Rare Earth Element (REE) mineralizations are not so “rare” in Sweden. They normallyoccur associated and hosted within granitic crystalline bedrock, and in mineral deposits together
with other base and trace metals. Major REE-bearing mineral deposit types are the apatite-ironoxide mineralizations in Norrbotten (e.g., Kiruna) and Bergslagen (e.g., Grängesberg) ore regions,the various skarn deposits in Bergslagen (e.g., Riddarhyttan-Norberg belt), hydrothermal deposits(e.g., Olserum, Bastnäs) and alkaline-carbonatite intrusions such as the Norra Kärr complex and Alnö.
In this study, analytical data of samples collected from REE mineralizations during the EURARE
project are compared with bedrock and till REE geochemistry, both sourced from databases available
at the Geological Survey of Sweden. The positive correlation between REE composition in the threegeochemical data groups allows better understanding of REE distribution in Sweden, their regional
discrimination, and genetic classiﬁcation. Data provides complementary information about correlationof LREE and HREE in till with REE content in bedrock and mineralization. Application of principal
component analysis enables classiﬁcation of REE mineralizations in relation to their host. These results
are useful in the assessment of REE mineral potential in areas where REE mineralizations are poorly
explored or even undiscovered.
Keywords: rare earth elements; Sweden; lithogeochemistry; till geochemistry

1. Introduction
The Rare Earth Element (REE) occurrences in Sweden are widely distributed all over the country
but some regions are more endowed than others. Major REE-bearing mineral deposit types, foundmainly in Norrbotten and Bergslagen ore regions, are the apatite-iron oxide mineralizations in
Kiruna and Grängesberg, the skarn mineralizations of the Riddarhyttan-Norberg belt, the alkaline
igneous rock-associated in Norra Kärr, and the hydrothermal mineralizations in Olserum and Alnö
deposit areas [1] (see Figure 1). Distribution of REEs in Swedish soil and bedrock has previously
been investigated by Sadeghi and Andersson [2] with the aim to identify the main changes in REE
geochemistry related to geology and weathering, taking also into account the current baseline levelfor REE in soil and bedrock over Sweden. The Geochemical Atlas of Sweden provides a harmonized,
countrywide database with modern baseline geochemical data from C horizon (element concentrationin the weathered parent rock horizon of soil proﬁle) in till [3] and REEs in diﬀerent solid media (topsoil,
subsoil, and stream sediments) have been investigated using the Forum of European Geological
Surveys (FOREGS) database in order to identify the REEs regional background values [4]. In Sadeghi
et al. [5], the Geochemical Mapping of Agricultural and Grazing Land Soil (GEMAS) project data were
investigated with focus on REEs in two solid media (topsoil from agricultural (Ap) and grazing land(Gr) soil) to identify the background values of REEs both in Sweden and in Europe. The Ap samples
Minerals 2020, 10, 365; doi:10.3390/min10040365

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Minerals 2020, 10, 365

were collected from regularly plowed ﬁelds at a depth interval of 0 to 20 cm and the Gr samples werecollected from soil under permanent grass cover at a 0 to 10 cm depth range.

Figure 1. Simpliﬁed map of major lithotectonic units and ore districts in Sweden, with a selectionof Rare Earth Elements (REEs) deposits, prospects, and occurrences (modiﬁed from [1] and bedrock
database at Geological Survey of Sweden)

REE deposits can be referred to regolith, basinal, metamorphic and magmatic associations basedon a mineral-systems approach [6]. Various deposit types form either directly from the crystallizationof the melt and/or ﬂuids predominantly derived from the melt. The magmatic deposits can be divided
into orthomagmatic and hydrothermal types. Basinal deposit types are inferred to be formed throughmechanical (e.g., placer) and chemical (e.g., phosphorite) sedimentary processes, and from diagenetic
ﬂuids generated in sedimentary basins. Deposit types of the regolith association require an REE-bearing
rock source to form feasible secondary contents of REE. The REE deposits are formed either due toenrichment of REE in the residual material and/or from local remobilization of REE. Deposit types ofthe metamorphic association are generated during regional and/or contact metamorphism and involve
related metamorphic ﬂuids.
This study analyzes, explores and evaluates the composition of REE mineralized samples collectedduring the EURARE project (2013–2017) enabling their genetic classiﬁcation into speciﬁc deposit types.
This information is integrated with the Geological Survey of Sweden (SGU)’s lithogeochemistry and tillgeochemistry databases to better approach and interpret the distribution of REE anomalous provincesat the scale of the country.

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2. Overview of REE Mineralizations in Sweden
2.1. REE Mineralizations in Granitic Pegmatites and Granitoids
Partial melting of crustal material produces felsic melts that are enriched in REEs. Elements inpegmatites that can be enriched up to minable (ore) grades include Li, Cs, Be, Sn, Nb, Ta, U, Y, Zr, andREEs. This long and diverse list of elements, to which a signiﬁcantly high content of B, P, F, Rb, Bi, Hf,
etc., could be added, make fractionated pegmatites among the most mineralogically complex deposit
types on the Earth [7]. Pegmatitic facies associated with syenite-alkaline granite complexes constituteimportant deposits of REEs, U, Nb and Zr, and less commonly, Be and P [7].
In Sweden, granitic pegmatites, typically of a moderately to highly fractionated character, often
host variable amounts of minerals rich in trace elements, including Be, Li, Nb, Ta, Sn, U, Th, as wellas REEs. [8,9]. One notable granitic pegmatite, that shows locally relatively high REE content,(a porphyritic granite to gneiss with age 1300 Ga so-called “RA-granite”) is located in Balltorp(Eastern Segment lithotectonic unit) (Figure 1) in southwestern Sweden [10]. The RA-granite is a
Be-F-Nb-REE-Sn-Ta-Th-U-Zr-anomalous gneissic granite that does show some potential but has so far
not been systematically explored (cf. [10] and references therein). Moreover, several other granitic tosyenitic rocks in southern Sweden and elsewhere exhibit increased REE contents. The most important
granitic pegmatites and granitoids with potential for REE mineralization are, Tåresåive (highest
contents of REEtot = 94,900 ppm) in northern Sweden (Norrbotten lithotectonic unit), Näverån inCentral Sweden (close to the Caledonides), as well as granitoids in Bergslagen (e.g., Ytterby; highest
contents of REEtot = 30,624 ppm), and granite in the southwestern part of Sweden, (e.g., Balltorp;
highest contents of REEtot = 8800 ppm), shown in Figure 1. A contrasting case to these primaryenrichments is the granitoid-hosted but epigenetic, shear zone-related mineralization at Näverån
(highest contents of REEtot = 6929 ppm) in central Sweden [1].
2.2. REE Mineralization in Alkaline Intrusive Rocks and Carbonatites
Several REE deposits associated with peralkaline complexes (syenitic rocks) are usually enrichedin U and Th and have a relatively high Heavy Rare Earth Element (HREE) content when comparedwith carbonatites [11]. In Sweden, Norra Kärr deposit represents one of the most advanced andpromising REE mining projects in northern Europe. This concentrically zoned alkaline intrusion, with
nepheline-syenite, contains eudialyte-group minerals highly enriched in REEs with up to 8.5 wt% (La +Ca + Nd + Y) [1], showing also a strong resource potential in heavy REEs.
Carbonatites are common hosts of orthomagmatic-type REE mineralizations that tend to bevariably enriched in REEs, Sr, Ba, U, Th, Nb, Ta, P, and F. Although the overall REE content ofcarbonatites may vary, their shape on the chondrite-normalized plot almost invariably displays highLight Rare Earth Element (LREE) content and no negative Eu anomalies [12]. Barium and Sr are
generally abundant, whereas U, Th, Nb, and P, are variably abundant in the mineralized carbonatite.
REE-bearing mineral phases have the highest REE content but major carbonate minerals, including
calcite and dolomite, may also contain substantial contents of REEs [13]. Sövitic carbonatitic melts
tend to show the highest levels of REE enrichment. They also show signiﬁcant enrichment in LREE(LREE/HREE ratio of ~40, compared to ~7 of alkaline and felsic melts).
In Sweden, the Alnö complex consists of alkaline and carbonatite intrusive rocks located on theisland of Alnö, in the Bothnian Bay and on the mainland north to northwest of Alnö, all along thecoast of east-central Sweden [14–16]. The carbonatite, and speciﬁcally the calcite-dominated søvite,
have signiﬁcantly elevated total REE contents, between ca 500 and 1500 ppm [17]. Alnö intrusivecomplex and associated co-magmatic rocks on the mainland have also been the focus of exploration
for REEs, including that carried out by the Boliden company in the 1970s. The main lithology ijolite
contains clinopyroxene with LREEs up to 100 times chondritic values [18,19]. The Söråker intrusion
located on the mainland north of Alnö comprises Ca- and Mg-carbonatite dikes and melilitolites.
Another small carbonatite occurrence is known from the Middle Allochthon in the Caledonides

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where calc-silicate rocks at Prästrun in west-central Sweden host a REE-Nb-(U-Th) mineralization(highest contents of REEtot = 2300 ppm). This intrusive-hosted REE mineralization was ﬁrst described
at Prästrun.
Alkaline lamprophyre is another lithology that tends to have elevated REE contents. Chemically,
potassic lamprophyre dikes are often characterized by relatively low SiO2 and TiO2, and high MgO
and K2 O contents. In addition, they show high content of large-ion lithophile trace elements (e.g., Rb,
Sr, and Ba) and LREE, but low content of high-ﬁeld-strength elements (e.g., Nb, Ta, Zr, Hf, andTi) [20]. An example of the REE-enriched lamprophyre is the 1.1 Ga alkaline ultramaﬁc dykes up to
1-metre-wide, outcropping along the northernmost part of the coast in the Gulf of Bothnia, (Kalix)
known as Storön deposit (Figure 1; [20,21]). The dikes represent olivine- and mica-rich lamprophyres
and silico-carbonatites [21].
2.3. REE Mineralization in Apatite-Iron Oxide Deposits
At present, over forty iron oxide apatite deposits are known in northern Sweden [22–24]. The main(at over 2000 Mt) deposit in the Norrbotten district is the famous Kiruna (Kiirunavaara) deposit andactive mine (cf. [24]). A hydrothermal (including replacement) and (ortho-) magmatic origin of ironoxide apatite deposits in Norrbotten have been suggested in numerous studies (e.g., [25–31] andreferences therein). The main REE host minerals in this deposit are apatite, typically a ﬂuorapatite,and monazite, occurring often as inclusions in apatite [32–35]. Apatite abundance and compositionin the ore varies, e.g., at Kiruna, where the phosphorus content in the upper parts of the ore is ca.
2%, equivalent to almost 11% apatite. REE enrichment is represented mainly by LREEs, Ce and Y
anomalies observed as early as the 1930s by Geijer [34]. REE geochemistry of the host rocks in thenorthern district shows two patterns, one with a negative Eu anomaly and strong enrichment in LREEs
at Kiruna and area close to this mine and the second one without a well pronounced Eu anomalye.g., at the Malmberget and the areas close to this mine (Figure 2) [36].

Figure 2. REE spider plot [37] of host rocks from the northern districts; Kiruna mine and occurrences close
to this area (red) display a distinct Eu trough and a relative enrichment in LREEs compared to the HREEs,
and while the samples from the Malmberget mine and occurrences nearby show extensive metamorphismand deformation overprint compared to the Kiruna area (Blue). Data sources: SGU and [36].

114Minerals 2020, 10, 365

In central Sweden, (the Bergslagen district), a number of iron oxide apatite deposits are knownwith the biggest being Grängesberg and Blötberget, which were mined in historical times (Grängesberg
mine closed in 1989) and both represent Kiruna-type deposits. The previous studies reveal that they
still constitute a rather important potential reserve of REEs and phosphorus, both from the main ore
and from mining waste [38]. The deposits are predominantly hosted by 1.9 Ga volcanosedimentaryrocks variably metamorphosed and deformed. The ore is usually enriched in LREEs and to a minor
degree in U, Th, Sm, Tb, Yb, and Y. The majority of REE enrichment is located in phosphate mineralssuch as apatite, xenotime, and monazite and to a lesser degree in silicates such as allanite andepidote [1,30,33,35].
2.4. Fe-REE Hydrothermal (Bastnäs-Type) Mineralizations
Evaluation of the REE geochemistry of hydrothermal mineral deposits, and the processes bywhich they are concentrated, is a complex task [39]. However, these mineralizations can be groupedon the basis of their geochemical and mineralogical characteristics. The degree of REE enrichment ina deposit is a function of the concentration of REEs in the ﬂuid, the water-rock ratio, the eﬃciencyof the precipitation process and the nature and amount of co-precipitated phases [39]. Oreskes andEinaudi [40] suggested that the high abundance of ﬂuorocarbonates and the lack of Ca minerals in
REE deposits (e.g., Olympic Dam) indicated that the ﬂuids behind hydrothermal REE mineralizations
may have been F- and CO2 -rich and that the REEs were transported as Cl and/or F complexes.
Lottermoser [41] suggested that the association of REE with U minerals indicated that U and REE werecomplexed by the same ligands and inferred that those were CO3 , F, or SO4 .
The Bastnäs-type Fe-REE deposits in Sweden are early Proterozoic, skarn-hosted iron oxide(magnetite-dominated), locally polymetallic (±Cu, Au, Co, Bi, Mo) mineralizations that in part carryvery REE-rich mineral assemblages (cf. [42]). They are located in the Bergslagen district and arecharacterized by the occurrence of locally abundant REE-rich silicate minerals such as cerite-(Ce) andallanite (sensu lato) but also include REE ﬂuorocarbonates such as bastnäsite-(Ce).
3. Materials and Methods
In this study the lithogeochemical and till geochemistry databases have been used for
re-interpretation of the REE baseline in diﬀerent rock types and till in Sweden, and for the evaluationof REE mineral potential. More than 20,000 rock samples were collected and analyzed as part ofthe bedrock mapping program at the Geological Survey of Sweden (SGU; www.sgu.se). Within the
EURARE project (www.eurare.org), ﬁeld work activities were carried out in several REE-mineralized
areas and were followed up by detailed investigations on selected mineralizations. The modal mineralcomposition, petrographical features and chemical compositions carried out for more than 200 rocksamples [1]. The lithogeochemical analyses were carried out at the ACME Lab (Vancouver, BC, Canada)
and ALS Scandinavia AB (Luleå, Sweden). At the ACME Lab the samples were analyzed for major,
minor, and trace elements by ICP-emission spectrometry following a lithium metaborate/tetraborate
fusion and dilute nitric digestion. At the ALS, trace elements including the full rare earth element
suites were obtained from fused beads followed by acid digestion and measured by either ICP-AES oran inductively coupled plasma sector ﬁeld mass spectrometer (ICP-SFMS).
The till database used in this study contains over 2500 samples analyzed by aqua regia digestion
followed by ICP-MS. These national-wide results have been published as the Geochemical Atlas of
Sweden [3] and are available for the public online (https://www.sgu.se/mineralnaring/geokemiskkartlaggning/geokemisk-atlas/).
The open-source software Geochemical Data Toolkit in R (GCDKIT-version 4.1) [43] which is built
using the freeware R language, has been applied for the data processing and biplot diagrams in thispaper (http://www.gcdkit.org).
Principal component (PC) analysis is a conventional multivariate technique that is often used forstudying geochemical data [4,44–47]. PC analysis reduces a large number of variables to a smaller

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number, allowing the user to determine the components (groups of variables) that account for variationin multivariate data [47]. PC analysis has often been used to process and interpret geochemical andother types of spatial data e.g., [4,5,47,48]. PC analysis builds on the correlation (covariance) matrix,
which measures the interrelationships among multiple variables. The ﬁrst PC (PC1) explains most of
the variance within the original data, and each subsequent PC (PC2–n) explains progressively lessof the variance. A multivariate dataset can usually be reduced to two or three PCs that account forthe majority of the variance within the dataset. PC analysis was performed on average values of thediﬀerent mineralization types as discussed above. Thirty-seven elements were chosen for PC analysis
based on correlation coeﬃcients. The Statistica software (version 13.2, Dell Software, Aliso Viejo, CA,
USA) has been used for the principal component analysis (https://statistica.software.informer.com/).
The Geochemical Atlas of Sweden [3] is a country-wide harmonized database with modern
baseline geochemical data from the C horizon in till. The database encompasses 2578 till samples from
SGU archive as well as new sampled till collected mainly in the mountainous areas of western Sweden.
Since the C horizon is considered as an anthropogenically undisturbed layer, the major geochemicalsignature in till should originate from underlying bedrock, its lithology, mineralogy, and potential
mineralization. Secondary processes such as ice transport, leaching, and biological activity have minor
impact on till chemical composition. During weathering, REEs are generally not very mobile, but thisvaries depending on the host mineral and local pH. REE mobility is further controlled by adsorption
onto iron oxides, phosphates, and clay minerals [2,49]. In general, till can be used as proxy for theunderlying bedrock in areas where outcrops are not available. Till can be treated then as composite
rock material and its composition can be often interpolated to the average parent rock.
4. Results and Discussion
4.1. REE Content in Bedrock
Lithogeochemistry data from the SGU database have been used to examine links between REE
contents in bedrock with known REE mineralizations. The main intention was to establish a tool
for future exploration and predictive mapping surveys. Generally, primary REEs are associatedwith magmatic (granite and pegmatite) and alkaline rocks. Rocks enriched in monazite and zircon
often contain higher contents of rare earth elements. In sedimentary rocks, higher contents of rareearth elements occur in shale and greywacke. There are no obvious diﬀerences in the geochemical
distribution between the LREE group (La, Ce, Pr, Nd, and Sm) and the HREE group (Eu-Lu + Y) in thelithogeochemical data. However, their relative contents may diﬀer for each element, probably due tothe mineral composition and type of the host rock. For example, Eu is often found in rock-forming
minerals such as plagioclase and in accessory minerals, mainly in allanite, bastnäsite, monazite, apatite,
zircon, and ﬂuorite. Typically, Eu can replace Sr and Ca in rock-forming minerals, therefore, itsenrichment in plagioclase is common. This results in Eu enrichment in plagioclase-rich rocks innorthernmost Sweden, such as gabbro and granodiorite.
The data with a higher content of the total REE and Y (above 500 ppm) were extracted from thedatabase and compared with the extracted data with a content of Y more than 100 ppm (Figure 3).
Our results show a linear correlation and match between the samples with a higher content of REE + Y
in total with Y content, and more importantly, there is a direct link between Y content in samples withexisting known mineralizations in Sweden (Figure 3).
4.2. Geochemistry of REE Mineralization in Sweden Inferred from the EURARE Project
During the EURARE ﬁeld campaign, approximately 200 samples were taken for petrography
and mineralogy studies and for geochemical analysis. The location of the samples and type ofmineralization are shown in Figure 4, and the geochemical results are provided in Table 1.

116Minerals 2020, 10, 365

Figure 3. Simpliﬁed map of lithogeochemical data in Sweden: (a) total REE + Y > 500 ppm and (b) Y
content >100 ppm (source of data: lithogeochemical database at the Geological survey of Sweden).

Figure 4. Simpliﬁed location map of samples in the EURARE project: (a) an overview of sample
locations and types, excluding magnetite-calc silicate skarn and BIF ones; (b) magnetite-hematite-quartz
samples; (c) magnetite-calc silicate skarn samples.

117

SampleKES155002A
PNY130005
ING090076A
BAST140001
PNY130001
Gyttorp
MÖRK2
JSM150007A
KES155003A
YBY140002
Rödbergsgruvan A
PNY130015
JSM150008A
HOLM140001
YBY140003
JSM150006AMÖRK3
GRUVHAG1
KES155006A
YBY140001
REU140001
SGUR10007
KES155001A
FLA140001
PNY130013
SVEAFALLEN
MÖRK1Occurrence/Deposit

Ingridstorp
StålklockanTåresåive
Bastnäs
Bastnäs
GyttorpMörkens
Djupedalsgruvan
Ingridstorp
Ytterby
RödbergsgruvanÖstanmossa
Djupedalsgruvan
Holmtjärn
Ytterby
Djupedalsgruvan
Mörkens
GruvhagenIngelsbo
Ytterby
Reunavaare
TybbleIngridstorp
Flakaberget
Johanna
Sveafallen
Mörkens

6,392,246
6,635,349
7,421,196
6,634,470
6,634,425
6,597,427
6,662,160
6,425,442
6,358,023
6,592,332
6,597,194
6,660,543
6,425,442
6,691,407
6,592,332
6,425,442
6,662,160
6,662,377
6,364,317
6,592,332
7,342,033
6,517,096
6,392,246
7,389,427
6,658,660
6,563,279
6,662,160

N
559,703532,852
709,317532,926
533,003497,660
577,465
578,440
372,718
690,226494,237
551,764
578,440
514,383
690,226
578,440
577,465
577,026
411,300
690,226
697,981
516,016
559,703
703,658
550,883
467,518
577,465

E
Pegmatite/aplite
Magnetite-Calc silicate skarn
Pegmatite/aplite
Magnetite-Calc silicate skarn
Magnetite-Calc silicate skarn
Magnetite-Calc silicate skarn
Magnetite-Calc silicate skarn
Fe-ore unspec.
Pegmatite/aplite
Pegmatite/aplite
Magnetite-Calc silicate skarn
Magnetite-Calc silicate skarn
Fe-ore unspec.
Pegmatite/aplite
Pegmatite/aplite
Fe-ore unspec.
Magnetite-(Fe-sulphide)-Calc silicate skarn
Magnetite-Calc silicate skarn
Pegmatite/aplite
Pegmatite/aplite
Pegmatite/aplite
Iron oxide apatite
Pegmatite/aplite
Pegmatite/aplite
Magnetite-Calc silicate skarn
Pegmatite/aplite
Magnetite-(Fe-sulphide)-Calc silicate skarn

Type47,17245,51242,651
38,07837,90737,06936,921
34,64333,89530,62430,53630,31929,99029,55026,41925,904
24,971
24,57422,192
18,695
18,207
17,42816,32015,556
14,00813,22011,077

TotREE(ppm)
43,00734,00133,05633,14933,13433,23333,22522,746
19,577
13,62429,37928,85016,626
12,550
10,239
11,133
21,636
23,6185192
3801
13,15817,274
16,0261627
12,966
12,579
9203

LREE
(ppm)
4165
11,511
9596
4929
4773
3836
3697
11,89814,31817,0001158
1469
13,36517,00016,18014,7713335
956
17,00014,8945049
154
294
13,929
1042
641
1874

HREE
(ppm)
U (ppm)
266
58.5
1505
66.4
92
1120
18.85
9
1330
1185
4.15
3.7
8.39
>10,000
1270
21.5
2.67
0.77
>10,000
2360
438
457
53.9
1600
8.68
16
2.47

Th
(ppm)
>1000
2.01
>1000
20.7
9.64
50.4
3.36
158
>1000
>1000
1.53
1.21
214
>1000
>1000
134
0.43
3.34
>1000
>1000
>1000
6.41
257
>1000
0.72
>1000
0.09

Table 1. Summary of samples analyzed with a total REE (TotREE) content above 1% (10,000 ppm). N and E are SwerefTM coordinates.

70
90
140
70
60
<10
40
2640
1260
6200
80
40
3360
450
1470
690
30
10
<10
6280
30
1400
90
520
30
1460
40

P (ppm)Minerals 2020, 10, 365118

Minerals 2020, 10, 365

The geochemical analyses of the samples collected during the EURARE project show that
68 samples have a total REE content above 1000 ppm and of those 27 have more than 1% total REE.
The samples with more than 1% REEs represent mineralizations classiﬁed as: iron oxide apatite(1 sample), Fe-ore unspeciﬁed (3 samples), magnetite-calc-silicate skarn (9 samples), magnetite(Fe-sulphide)-Calc silicate skarn (2 samples), and granitic pegmatite (12 samples). These results are
summarised in Table 1.
A bivariate plot of total REE content vs. LREE (Figure 5a) and vs. HREE (Figure 5b) show that
there is a general bimodal linear correlation between total REE and LREE (Figure 5a), one trend ischaracterized by LREE enrichment (for magnetite-skarn mineralization) and a second one by LREE
depletion, especially in some late magmatic pegmatite REE mineralizations and some related to ironmineralizations. The comparison of total REEs with HREEs also shows clear decoupling between
two groups, the ﬁrst composed of low HREE mineralization types (e.g., various iron ores, skarn
mineralizations, carbonatite, and sulphide-dominated skarn ores) and these with elevated HREE
content are mainly represented by late magmatic pegmatites and unspeciﬁed iron ore types. It is
unclear how the magmatic rocks diﬀer in primary composition, so they host both low- and high-level
HREE mineralizations. BIF deposits (magnetite-hematite-quartz) do not show a very high content of
REEtot and generally are most likely to have elevated LREE rather than HREE. The number of two
samples of sulﬁde dominated skarn is limited and, therefore, it is excluded from interpretation of theresults to avoid any over interpretation.
A bivariate plot of La vs. La/Yb (Figure 6) shows a weak but overall positive correlation between
the two variables, except for a few samples, which show higher HREE contents. This suggests thatthose samples with a high LREE content also have relatively high LREE/HREE ratios. They mostly
represent REE mineralizations associated with granitic pegmatites. One sample from iron oxide-apatite
deposits contains REE-bearing minerals such as xenotime and monazite and shows higher contents of
HREEs (Figure 6).
Thorium and uranium content and their ratio are useful for recognizing geochemical facies
and also the content of radioactive elements in diﬀerent types of mineralization which may have an
impact on future exploitation. Figure 7a shows Th content versus total REEs. In general, skarn andiron oxide-apatite mineralizations with higher contents of total REEs (REEtot > 1%) show a positive
correlation with low contents of Th (Th < 50ppm). A few samples from iron oxide mineralizations
show higher content of Th up to 200 ppm and total REE contents of up to 3.5%. Two samples of granitic
pegmatites show higher contents of Th, whereas those of total REEs are low, which means that thecontent of Th in those samples is not directly related to REE-bearing minerals and may also correspondto the primary origin of the host rocks and its later deformation. For example, rocks with sedimentary
origin have usually higher Th/U ratios than magmatic rocks, and these ratios can be further modiﬁed
during regional metamorphism [11,12,49].

119

Minerals 2020, 10, 365

Figure 5. Bivariate diagrams of total REEs versus LREEs (a) and HREEs (b) for EURARE project samples.

120Minerals 2020, 10, 365

Figure 6. Bivariate plot of La (ppm) versus La/Yb for the EURARE samples (see the legend in Figure 5).

The U/total REE (Figure 7b) of the EURARE samples shows that high U contents (U > 400 ppm)
are typical of REE mineralization hosted in granitic pegmatites. An additional few samples collectedfrom magnetite-calc silicate skarn type deposits show higher contents of U (>1000 ppm). In the sample
from magnetite-calc silicate skarn, one sample shows a high content of total REE correlates with highU content which may indicate the U enrichment in REE-bearing minerals.
Comparison of total REE with Th/U ratios based on two analyzed samples from lamprophyres(Figure 7c) shows that the samples with low contents of total REE (REEtot < 400) typically have higher
Th than U contents (Th/U > 20). The Th/U ratio is generally low in the samples from skarn andiron oxide-apatite REE mineralizations (Th/U < 10) with the exception of four samples from ironoxide-apatite REE mineralizations (at the Djupedalsgruvan) that show high contents of both Th and
REE (and contain xenotime and allanite minerals) (Figure 7c).
Plots of quantitative values of selected REEs (Nd, Pr, and Y) and phosphorus compared to binary
plots of LREE versus HREE show at least two contradicting trends. For example, in Figure 8a samples
with high Nd contents show two opposite trends; one positive with well correlated LREE and Ndcontents accompanied by low HREE and the second negative trend where samples with decreasing
Nd contents plot towards a lower LREE content and a higher HREE. Few exceptions of samples withhigh content of HREE and Nd content (>4000 ppm) can also be observed (Figure 8a).

121

Minerals 2020, 10, 365

Figure 7. Cont.

122

Minerals 2020, 10, 365

Figure 7. Binary plots of (a) total REE vs. Th (ppm); (b) total REE vs. U (ppm); (c) total REE vs. Th/U;
for the EURARE samples (see the legend in Figure 5).

In a similar type of plot, quantitative content of Pr in the samples shows a semi-concordant trendwith Nd, and this may reﬂect that the minerals bearing Pr and Nd are the ones containing HREE,
such as xenotime-(HREE) and fergusonite-(HREE) as well as gadolinite-(REE).
Figure 8c shows linear positive correlation of high Y samples with a higher content of HREE.
Since Y has a similar geochemical behavior as HREE the low HREE contents follow low Y contents in
EURARE samples. High LREE samples are less aﬀected by Y content with weak negative correlation.
Samples with low and high Y content when compared to LREE might reﬂect various degrees ofhost magma diﬀerentiation or secondary remobilization during regional metamorphism (competitive
re-crystallization and crystallization of metamorphic minerals) [11,12,49].
Comparison of phosphorus content, indicative of the presence of phosphate minerals shows two
diﬀerent clusters (Figure 8d). The ﬁrst cluster shows that some high P samples have low contents
of both LREEs and HREEs, which can be interpreted as apatite mineralization in iron ore with noREE enrichment. The second cluster shows that samples with high P content correlate with highHREE content, which may indicate the presence of minerals other than apatite as a host of HREEs,
such as xenotime.
Within the sample set collected during the EURARE project, there are several iron ore samples(Fe-ore unspec.) that cannot be easily classiﬁed (more mineralogical and petrological investigations areneeded). In an attempt to classify them, the classical PM-normalized REE diagram has been used [37](McDonough and Sun 1995) (Figure 9). At a ﬁrst glance, their REE patterns suggest the presence oftwo genetic groups. The ﬁrst group with REE enrichment (>100 ppm) shows a similar trend to the ironoxide-apatite- and skarn-type REE mineralizations, with a negative Eu anomaly and relative HREE
enrichment. The second group shows a weak positive Eu anomaly and steep decreasing trends from
LREEs to HREEs accompanied by a rather low total of REEs (<100 ppm). These two samples may havean alkaline or peralkaline origin but a more detailed study is needed (Figure 9).

123Minerals 2020, 10, 365

Figure 8. Cont.

124

Minerals 2020, 10, 365

Figure 8. Plots of quantitative values of selected elements compared to LREEs and HREEs in the
EURARE samples: (a) Nd (ppm); (b) Pr (ppm); (c) Y (ppm); (d) P2 O5 (wt%). Samples with contents
less than background values are excluded.

125

Minerals 2020, 10, 365

Figure 9. Primitive mantle-normalized [37] REE diagram of iron oxides of unknown origin.

4.3. Principal Component (PC) Analysis
The results of PC analysis are reported in Table 2 for the PCs with eigenvalue >1 and show that
PC1 and PC2 explain 30% and 17%, respectively, of the total variance among the REE-mineralized
samples. It implies that in PC1 with an association of REEs, and Co, Fe, and Mn most likely represent
magnetite calc-silicate skarn samples, while an association of Cs-Hf-Nb-Sr-Ta-Th-Zr-Al-Ba in PC2
likely indicates REE mineralizations related to granitic pegmatite bearing Nb-Ta mineralization.
Table 2. Principal component loading, eigenvalues, % variance explained, and cumulative % variancefor PC1 to PC8.
Eigenvalues of Correlation Matrix and Related Statistics
Value Number
PC1
PC2
PC3
PC4
PC5
PC6
PC7
PC8

Eigenvalue% Total (Variance)

Cumulative %(Eigenvalue)

Cumulative %Variance

17.56
9.99
4.65
3.41
2.89
2.20
1.94
1.86

29.77
16.93
7.88
5.78
4.90
3.74
3.30
3.16

17.56
27.55
32.20
35.61
38.50
40.70
42.64
44.51

29.77
46.69
54.57
60.35
65.25
68.98
72.27
75.43

The PC3 deﬁnes an association of Cs-Rb-K-Li, which probably represents alkaline and/or felsic
rocks including granitic pegmatite and related polymetallic mineralization (Table 3).

126Minerals 2020, 10, 365

Table 3. Explanation of the ﬁve principal components for elements analyzed, accounting for 65% of thetotal variance.
Centered Log-Ratio (clr)-Transformed Data

Component% of Variance
Explained

Association

Interpretation

PC1

29.77

(i) REEs(ii) Fe-Mn(i) REEs associations(ii) Magnetite calc-silicate skarn

PC2

16.93

(i) REEs(i) REEs associations(ii) Granitic pegmatit-bearing
Nb-Ta mineralization(ii) Cs-Hf-Nb-Sr-Ta-Th-Zr-Al-Ba
PC3

7.88

(i) Cs-Rb-K-Li(ii) Ag-Cd-Bi-Co-Cu-Zn-C(i) alkaline/granitic rocks(ii) Base metals component

PC4

5.78

(i) C-Hf-Sr-Ba-Ca(ii) Fe- -Mg-Tl(i) Carbonatite/lamprophyr?
(ii) Magnetite calc-silicate skarn?

PC5

4.9

(i) Hf-Zr-In(ii) V-P-Ni-Cr(i) Felsic component(ii) Maﬁc component

Plots of pairs of PCs show that it is likely possible to distinguish between some diﬀerent types of
REE mineralization or to distinguish between samples with a higher content of REEs and those withlow one (Figure 10). For example, in the plot of PC1 versus PC2 (Figure 10) there are several clusters
representing an association of REEs and granitic pegmatite bearing Nb-Ta mineralization. A plot of
PC3 versus PC4 distinguishes between REE mineralization associated with magnetite calc-silicate
skarn, alkaline, and/or granitic rocks and polymetallic mineralization (Figure 10).
4.4. Comparison of REE Till Geochemistry with Lithogeochemistry of Underlying Bedrock
Table 4 shows simple statistics of REE distribution in till. The median values are similar to upper
crust values by Rudnick and Gao [49]. Due to the fact that Rudnick and Gao’s values are calculated for
total contents and till results originate from aqua regia digestion, one of the conclusions is that theSwedish till is enriched in REEs in comparison to the global values.
Table 4. Percentiles, minimum, and maximum values for the REE + Y + Sc from the till dataset. P50 isa median value. * abundance in the upper continental crust [49].
Element

n = 2578

UCC *

Min

P25

P50

P75

P90

P99

Max

La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu

ppm
ppm
ppm
ppm
ppm
ppm
ppm
ppm
ppm
ppm
ppm
ppm
ppm
ppm

31
63
7.1
27
4.7
1.0
4.0
0.7
3.9
0.8
2.3
0.3
2.0
0.3

8.1
16.3
1.9
7.0
1.4
0.1
1.1
0.2
0.9
0.1
0.4
0.1
0.3
0.0

25.6
60.1
6.3
23.3
4.4
0.6
3.6
0.5
2.9
0.6
1.5
0.2
1.3
0.2

31.6
74.8
7.7
28.5
5.4
0.8
4.4
0.6
3.5
0.7
1.9
0.3
1.7
0.2

39.6
96.0
9.4
34.9
6.7
1.0
5.4
0.8
4.3
0.9
2.3
0.3
2.1
0.3

49.9
123.7
11.8
43.4
8.3
1.2
6.6
1.0
5.3
1.1
3.0
0.4
2.7
0.4

85.5
211.8
20.6
73.1
13.3
1.9
10.4
1.5
8.5
1.7
4.6
0.6
4.2
0.6

198.9
388.1
38.1
129.3
24.2
5.5
22.3
3.6
22.6
4.8
14.3
2.0
14.6
2.6

Y
Sc

ppm
ppm

21
14

3.7
0.6

14.9
3.5

18.3
4.5

22.6
5.6

27.9
6.8

47.1
9.6

163.2
21.2

127

Minerals 2020, 10, 365

Figure 10. Biplots of (a) PC1 vs. PC2; (b) PC3 vs. PC4.

128

Minerals 2020, 10, 365

Maps of LREE and HREE contents in till (see [3]) show both similar and distinct distribution
patterns. Generally, LREE (La-Sm) and HREE groups (Gd-Lu, Y) have similar distribution within
the group, therefore, we use here the La till map as a representative for all LREEs and the Yttrium
map as a representative for HREEs (Figure 11). Raster La and Y maps with regional anomalies show
large-scale REE trends, while plotted lithogeochemistry results (circle symbols) underlie local spots
and extreme values.

Figure 11. Geochemical maps comparing till geochemistry (raster, after Andersson et al., 2014 [3]) withlithogeochemistry (circle symbols) in Sweden: (a) Y in till and bedrock; (b) La in till and bedrock.

LREEs and Eu have major anomalies in northernmost Sweden (outlying Archean lithologies ofmagmatic origin) while these anomalies are absent in HREEs. Minor enrichment in LREEs visible in129

Minerals 2020, 10, 365

lithogeochemistry data reﬂects possible LREE enrichment in granite-pegmatite bedrock and apatite-irondeposits near Kiruna.
Further south, in central Norrbotten (Figures 1 and 11), the bedrock (granite-pegmatite and felsic
metavolcanics rocks) and overlying till are highly enriched both in LREE and HREE and correlation
with many REE (+W, Mn, and U) mineralizations, e.g., Tåresåive (REEs, Mo, U, Th, Nb, and Ta).
In central Sweden, LREE and HREE enrichment in till occurs in the Caledonides within the tectonic
windows where many sulﬁde mineralizations and U deposits have been mapped. LREE enrichment at
the Caledonian front and within the Lower Allochthon can be attributed to the REE enrichment in theﬁne-grained sedimentary rocks (black shale and phosphorite), e.g., Tåsjö deposit (Figures 1 and 11).
On the Fennoscandian Shield side, few REE mineralizations are known and they possibly may bereﬂected in till with its prominent anomaly south of Östersund, possibly in relation to the so-called
Rätan granite. NE from Östersund, in central Sweden, -there is a prominent HREE anomaly and it isprobably related to the outcrops of younger granite (so-called Revsund granite). The lithogeochemistrydatabase has poor coverage in central Sweden, however, Alnö carbonatitic intrusion by the coastis both visible in till and few collected rock samples with elevated LREE contents. In the Ljusdal(Ljusdal lithotectonic unit, Figure 1), high HREE (and elevated LREE) contents can be noticed in tilland several small REE (W, Mn, U, and Fe-Mn) mineralizations are known in direct relation to theanomaly extension. Lithogeochemistry data are missing in this region.
In the Bergslagen ore district, there is a general decoupling between REE lithogeochemistry and
REE content in till. The known REE mineralization line in western Bergslagen is not visible in tillgeochemistry. Point REE till anomalies can be observed in the NW part of Bergslagen (often related toiron mineralizations) and LREE high contents in till occur in southern Bergslagen by the coast.
In southern Sweden, LREE anomalies in the bedrock are often related to mineralizations andgranitic lithologies. Younger granites of the Blekinge lithotectonic unit (Figure 1) and ﬂuorite
mineralization in southern Sweden also seem to be well reﬂected in till geochemistry. High contents of
Y in bedrock and till in southern Sweden, south and south-west of the lake Vättern (Figure 11) seem tocorrelate with occurrences of highly metamorphosed rocks including occurrences of eclogites with
Y-enriched garnet and pyroxene. Minor REE (W, Mn, and U) occurrences are known north of the Y
till anomaly.
The comparison of till geochemical maps with lithogeochemistry and mineralizations shows clear
correlation between REE anomalies in surﬁcial deposits and the underlying bedrock and the presence ofmineralization. Therefore, till can be eﬀectively used for REE mineral exploration and mineral potential
studies in areas where there is no suﬃcient information about the bedrock and mineralizations.
5. Conclusions
The analysis of REE geochemistry in mineralized samples compared to bedrock lithogeochemistry
and till geochemistry reveals several regional trends in REE distribution in Sweden, which may havepractical implications for mineral exploration:
1.
2.
3.

4.

Lithogeochemistry of the bedrock correlates with geochemical REE provinces and even single
REE ore deposits;
REE geochemistry of mineralized EURARE samples allows geochemical classiﬁcation of various
REE occurrences in Sweden;
Till geochemistry presents the complex picture of REE distribution at the country scale. Many
known REE locations are underlain by till anomalies. On the other hand, there are many high
REE spots in till, which cannot be easily explained by known REE deposits. They may presentinteresting exploration targets;
Taking into account favorable geological and metallogenetic conditions in Sweden, there is agood potential for discovering new REE deposits;

130Minerals 2020, 10, 365

5.

Applied geochemical methods should be integrated in a multidisciplinary way with complementary
data and methods, such as mineralogy, geophysics, and numerical modeling and statistics,
to improve understanding of REE ore genesis and related geo-modeling. This holistic approach,also including ore geochemistry and mineralogy, will provide a well-documented REE mineralknowledge base, intelligence, and exploration methodology.

Author Contributions: Conceptualization, M.S. and N.A.; methodology, M.S. and A.L.; software, M.S. and A.L.;
validation, M.S., A.L. and N.A. and.; formal analysis, M.S. and N.A.; investigation, M.S.; resources, M.S.; data
curation, M.S.; writing—original draft preparation, M.S and A.L.; writing—review and editing, N.A. and A.L.;
visualization, M.S and A.L. All authors have read and agreed to the published version of the manuscript.
Funding: The EURARE project was funded by the European Community Seventh Framework Programme(FP7/2007-2013) under grant agreement no. 309373. The project, number 35242, was co-funded by the Geological
Survey of Sweden (SGU).
Acknowledgments: We thank Erik Jonsson, Torbjörn Bergman, Magnus Ripa, Per Nysten, Johan Söderhielm, andDick Claeson for their contributions to the EURARE project. The EURARE project was funded by the European
Community Seventh Framework Programme (FP7/2007–2013) under grant agreement no. 309373. The project,number 35242, was co-funded by the Geological Survey of Sweden (SGU). The authors thank two anonymous
reviewers for their constructive and valuable comments and suggestions.
Conﬂicts of Interest: The authors declare no conﬂict of interest.

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minerals
ArticleEﬀect of Mineralogy on the Beneﬁciation of REE from
Heavy Mineral Sands: The Case of Nea Peramos,
Kavala, Northern Greece
Christina Stouraiti 1, *, Vassiliki Angelatou 2 , Soﬁa Petushok 1 , Konstantinos Soukis 1
and Demetrios Eliopoulos 2
12*Faculty of Geology and Geoenvironment, National and Kapodistrian University of Athens,
15784 Athens, Greece; sofpetushok@geol.uoa.gr (S.P.); soukis@geol.uoa.gr (K.S.)
Department of Mineral Processing, Institute of Geology and Mineral Exploration, 13677 Acharnes, Greece;
vasaggelatou@igme.gr (V.A.); deliopoulos1000@gmail.com (D.E.)
Correspondence: chstouraiti@geol.uoa.gr; Tel.: +30-210-727-4941

Received: 24 February 2020; Accepted: 22 April 2020; Published: 26 April 2020

Abstract: Beneﬁciation of a rare earth element (REE) ore from heavy mineral (HM) sands by particle
size classiﬁcation in conjunction with high-intensity magnetic separation (HIMS) was investigated.
The HM sands of Nea Peramos, Kavala, Northern Greece, contain high concentrations of REE
accommodated mainly in silicate minerals, such as allanite. However, the potential of the Northern
Greek placer for REE exploitation has not been fully evaluated due to limited on-shore and oﬀ-shore
exploration drilling data. Characterization of the magnetic separation fractions using XRD andbulk ICP-MS chemical analysis showed that the magnetic products at high intensities were strongly
enriched in the light REE (LREE), relative to the non-magnetic fraction. Allanite and titanite are themajor host mineral for REE in the magnetic products but mainly allanite controls the REE budget due
its high concentration in LREE. SEM/EDS and ICP-MS analysis of the diﬀerent particle size fractions
showed LREE enrichment in the fractions −0.425 + 0.212 mm, and a maximum enrichment in the−0.425 + 0.300 mm. The maximum enrichment is achieved after magnetic separation of the particle
size fractions. Mass balance calculations showed that the maximum REE recovery is achieved aftermagnetic separation of each particle size fraction separately, i.e., 92 wt.% La, 91 wt.% Ce, and 87 wt.%
Nd. This new information can contribute to the optimization of beneﬁciation process to be appliedfor REE recovery from HM black sands.
Keywords: rare earth elements (REE); heavy mineral sands; EURARE; allanite; monazite;
HIMS; mineralogical characterization; geochemical characterization; magnetic separation; particle
size fractions

1. Introduction
Heavy mineral sands (or black sands) are coastal deposits of resistant dense minerals that locallyform economic concentrations of the heavy minerals. They serve as a major source of titanium
worldwide with main minerals rutile and ilmenite and, in some cases, show high accumulation inrare earth elements (REE) and Th [1]. The rare earth elements (REE) are a group of 17 chemicallysimilar elements, the lanthanides, scandium (Sc), and yttrium (Y) which behave similarly in mostenvironments in the Earth’s crust. REE are considered as “critical metals” for the European Union
economy due to the vast application in a variety of sectors, a complicated production process as wellas political issues associated with the monopoly in supply from China, especially the supply of heavyrare earths [2–4]. According to the recent EU Joint Research Center (JRC) report among the REE, six are
identiﬁed as more critical, because their combined importance for strategic sectors of the economy

Minerals 2020, 10, 387; doi:10.3390/min10050387

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Minerals 2020, 10, 387

such as high-eﬃciency electronics and energy technologies with risks of supply shortage. These are
Dy, Eu, Tb, Y, Pr, and Nd [5].
Currently REE are not exploited in Europe, however, due to the current situation several
exploration projects have been assessed in the course of the recently ended EURARE and ASTER
European projects, which showed that some of them are in an advanced stage of exploration anddevelopment [6,7]. The most promising cases are the alkaline igneous rock-hosted deposits in SouthGreenland, the Norra Kärr deposit in Sweden and Fen Complex in Norway (Goodenough et al. [6]) andthe alkaline volcanic-derived placers of Aksu Diamas in Turkey [8]. In Greece, the most signiﬁcant REE
concentrations are associated with heavy mineral sands on the coast of Nea Peramos and Strymonikos
Gulf. Moreover, the EURARE project highlighted the signiﬁcance of the secondary REE deposits,
such as the bauxite residue (red mud) from the processing of Greek bauxites [4,6,9].
Mudd and Jowitt [1] stretched the economic potential of heavy mineral sands as an importantunderestimated REE resource especially for monazite and xenotime minerals. HM sands remain
excluded from mineral resource considerations mostly due to the environmental problems that areassociated with the radioactivity of tailings and the reagents used. However, there has been limited
research work on the quantiﬁcation of these impacts [3].
Geological prospecting by the Institute of Geology and Mineral Exploration (I.G.M.E.) of Greeceon the black sands in the broader area of Strymon bay started in 1980’s and focused primarily on
the natural enrichment in actinides (U-Th) and associated radioactivity in the on-shore and oﬀshore
zones of Loutra Eleftheron to Nea Peramos regions [10–13] (Figure 1). In the last decade, there isan increasing number of geochemical and mineralogical studies that have been carried out on the
coastal areas of Kavala [14–19], Sithonia Peninsula of Chalkidiki [20], Touzla Cape [21], and the area ofMaronia, Samothrace [22]. Most of these studies focused on the characterization of the placers andtheir natural radioactivity ([23] and references therein). Previous studies have demonstrated that theheavy minerals, monazite, allanite, titanite, uraninite, zircon, and apatite, are traced in the Kavala
black sands, derive from the Symvolon/Kavala pluton, a deformed granodioritic complex of Miocene
age (Table 1) [6,9,14,24]. Despite the low grade of the Greek placers at Nea Peramos, as emphasizedin the EURARE project, there is a good potential of beneﬁciation due to the coarse particle size andthe liberation of REE minerals. The Northern Greece heavy mineral sands potential was not feasible
to be fully evaluated as a potential REE resource in the course of EURARE project due to limited
exploration data.
Worldwide, the commonly exploited rare earth-bearing minerals in industrial scale are bastnäsite,monazite, and xenotime [25]. Other REE-bearing minerals such as eudialyte, synchysite, samarskite,allanite, zircon, steenstrupine, cheralite, rhabdophane, apatite, ﬂorencite, fergusonite, loparite,
perovskite, cerianite, and pyrochlore are rarely found in deposits of economic signiﬁcance [26].
However, there are new deposits being under development containing many new REE minerals that
seek further understanding, such as zircon, allanite, and fergusonite [27].
Beneﬁciation of the three commercially extracted heavy minerals, bastnäsite, monazite, andxenotime involves gravity, magnetic, electrostatic, and ﬂotation separation methods with froth ﬂotation
being the most commonly applied REE mineral separation operation ([27–29] and references therein).
There are numerous research articles on REE mineralogy and hydrometallurgical processing but there
is still a lack of comprehensive descriptions of the beneﬁciation methods necessary to concentrate REE
minerals. A main reason for this lack is the fact that concentrates of monazite and xenotime worldwide
are produced from heavy mineral sands, therefore, comminution is scarcely required [26].

136Minerals 2020, 10, 387

Figure 1. Geological map of the sampling area of Nea Peramos Loutra Eleftheron coast, Kavalaregion [12,13]. The large igneous body behind the coastline is Symvolon granite.
Table 1. Common silicate, phosphate, and carbonate rare earth element (REE) bearing minerals inheavy mineral (HM) sands (bold, this study) and bauxite residues in Greece [6,9,14,24]. Data for mineralproperties from [26] (and references therein) and apatite data from [30].
REE-Mineral

Chemical FormulaDensity(g/cm3 )Magnetic
Properties

Weight
% REO

ThO2

UO2

(Ce,Ca,Y)2 (Al,Fe2+ ,Fe3+ )3(SiO4 )3 (OH)
(Y,Ce,Ca)2 (Al,Fe3+ )3 (SiO4 )3 (OH)
(Ca,Ce,Th)(P,Si)O4(Ca,REE)TiSiO5(Th,U)SiO4(Zr,REE)SiO4

3.50–4.20
n/a
5.28
3.48–3.60
6.63–7.20
4.60–4.70

paramagnetic
paramagnetic
n/aparamagnetic
paramagnetic
diamagnetic

3–51
3–51
<3
<3
-

0–3
0–3
<30
70–80
0.1–0.8

10–16-

Ca5 (PO4 )3 (F,Cl,OH)(Ca,Ce)5 (PO4 )(Ce,La,Nd,Th)PO4(La,Ce,Nd,Th)PO4(Nd,Ce,La,Th)PO4(Ce,La)PO4 .H2 O
YPO4

3.17
3.10–3.25
4.98–5.43
5.17–5.27
5.43
3.77–4.01
4.40–5.10

n/an/aparamagnetic
paramagnetic
paramagnetic
n/aparamagnetic

~19
35–71
35–71
35–71
52–67

0–20
0–20
0–20
-

0–16
0–16
0–16
0–5

(Ce,La)(CO3 )F(La,Ce)(CO3 )F
Y(CO3 )F
Ca(Nd,La)(CO3 )2 F

4.9–5.2n/a
3.90–4.00
4.11 (calc)

paramagnetic
paramagnetic
paramagnetic
n/a

70–74
70–74
70–74
-

0–0.3
0–0.3
0–0.3
-

0.09
0.09
0.09
-

Silicates
Allanite (Ce)Allanite (Y)Cheralite (Ce)Sphene (titanite)
ThoriteZirconPhosphates
ApatiteFluorapatite
Monazite (Ce)Monazite (La)Monazite (Nd)
Rhabdophane (Ce)Xenotime (Y)Carbonates
Bastnäsite (Ce)Bastnäsite (La)
Bastnäsite (Y)Synchysite (Nd)n/a: not available; (-): this information is not known.

This paper presents results of an ongoing beneﬁciation study of REE-rich HM sands form the Nea
Peramos, Kavala (Northern Greece) at laboratory scale. Two process schemes were applied and tested
at laboratory scale in this stage in order to improve understanding of mineral separation by diﬀerent
physical processes including, particle size classiﬁcation through wet sieving as a preconcentration
137

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process followed by high-intensity magnetic separation (HIMS) [26,27]. Scanning electron microscopy(SEM), X-ray diﬀraction analysis, and inductively coupled plasma–mass spectrometry (ICP-MS) wereused in combination, for the characterization of feed material as well as beneﬁciation products.
Geological Setting
The studied area is situated in Northern Greece, in the internal domain of the Aegean arc (Figure 1).
It is part of the Rhodope massif, a polymetamorphosed nappe stack, which comprises high-grade
Pre-Alpine felsic to intermediate orthogneisses (mostly Variscan age), schists, and marble interlayeredor tectonically overlain by basic to ultrabasic rocks, that were exhumed in the Cenozoic [31–38].
The lowermost Pangaion-Pirin or Lower Unit is exposed in several domes in the Rhodope area
and as the footwall to the top to SW Strymon detachment [37,39]. The Pangaion-Pirin Unit
comprises a thick marble sequence alternating with schists, both underlain by Variscan orthogneiss
and schists [37,40,41]. The Pangaion-Pirin Unit is intruded by several Oligocene to early Miocene
granitoids [37,42]. The syn-tectonic early Miocene Symvolon (or Kavala) granodiorite, which occupies
a large part of the studied area, has intruded and deformed along the southeastern end of the Strymon
detachment [39,41]. Marine deposits of Pliocene age, mainly sandstones and marls, are observed along
the coast, covering the granodiorite.
2. Samples and Methods
2.1. Sampling
Fifteen samples of coastal sands were collected alongside a coastal line of 10 km from Loutra
Eleftheron to Nea Peramos areas, Kavala region, Northern Greece (Figure 1). Sample campaign was
carried out in the course of EURARE project during the period of 2013–2015. The sampling techniqueinvolved opening holes of 40 cm deep. According to ﬁeld observations, locally, there were black sandconcentrations in layers of few cm thick in alternation with typical light-colored sands. Due to sample
inhomogeneity a large amount of ~15 kg of bulk sand from each sampling location was taken for
securing a representative sample. A 20 kg sample of placer sand with an initial top size of 1.7 mm was
prepared by mixing equal weights of individual collected samples (composite A-mixed sample) anddivided by splitting into samples of 2.5 kg.
2.2. Sample Characterization
2.2.1. XRD Analysis
The major mineralogy was determined by powder X-ray diﬀraction using a Siemens D-5005
diﬀractometer with Cu K radiation, at the Geology Department of NKUA. Intensities were recorded
at 0.02◦ 2θ step intervals from 3◦ to 70◦ , with a 2 s counting time per step. The resultant diﬀraction
patterns were processed using EVA software by Bruker AXS Inc (Madison, WI, USA), in order ﬁrst to
identify peaks and then relate them to selected mineral phases that are present in the Kavala blacksands samples (Figure 2). Detection limits are of the order of 1 wt.% approximately, but this is mineraland sample dependent.
The magnetic and non-magnetic fractions and all the particle size fractions were ﬁnely ground in
the appropriate size, placed in a sample holder, and smeared uniformly onto a glass slide, assuring a
ﬂat upper surface.
Semiquantitative analysis was performed using the Reference Intensity Ratio method (RIR) of
I/Ic [43]. For the semiquantitative analysis, the Diﬀract Plus software by Bruker-AXS was used.
The minerals that did not have I/Ic ratios, the ratios from similar minerals of the same mineral groupwere used. For allanite, intensity ratio of epidote (0.9) was used, and for magnesiohornblende, anaverage ratio of hornblende (0.6) was used.

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2.2.2. SEM/EDS Analysis
Sand particles were mounted within a plug of epoxy resin and then polished particle-mountsections were prepared. Before SEM analysis, the thin section was covered with a thin veneer of carbon
using a vacuum carbon coater. Textural analysis and semiquantitative elemental analysis of heavy
minerals was undertaken using a JEOL JSM-5600 Scanning Electron Microscope (SEM) coupled to anenergy dispersive X-ray spectrometer (EDS) of OXFORD LINK ISIS 300 (OXFORD INTRUMENTS),
with the use of software for ZAF correction, at the Faculty of Geology and Geoenvironment, NKUA,
using secondary electron (SE) and backscatter electron (BSE) modes.
2.2.3. Bulk Chemical Analysis
REE ore sample preparation included initial digestion with Aqua Regia and HF (in order to solve
dissolution issues in silicate samples) in Teﬂon® containers. Subsequently, the residue was chemicallyattacked with HCl and H2 O2 . Finally, the samples are preserved in 5% concentrated HCl. REE wereanalyzed by inductively coupled plasma-mass spectrometry (ICP-MS) at I.G.M.E. (Supplementary
Table S1). External calibration solutions with matrix correction were used to measure the instrumental
sensitivity of ICP-MS.
2.3. Beneﬁciation Tests
The beneﬁciation tests were conducted at the Laboratory of Beneﬁciation and Metallurgy, I.G.M.E.
2.3.1. Particle Size Analysis
From the composite sample A-mixed, a representative sample (of 2327 g initial weight) was
prepared properly in order to be submitted to particle size analysis. Samples were submitted to wetsieving in order to provide data on particle size distribution, using sieves with aperture of 1.70, 0.850,
0.500, 0.425, 0.355, 0.300, 0.212, and 0.150 mm. All the samples were placed in a dryer for 24 h at 90 ◦ C.
The information acquired by particle size distribution contributes to the study of the eﬀects ofparticle size, mineral liberation, and association characteristics aiming at the selection of appropriate
process schemes to be tested and used. In the course of wet sieving process, a change in the color wasobserved as we moved onto smaller particle size fractions.
Preconcentration step helps in rejecting early gangue materials thus achieving beneﬁts such as
higher feed grades, lower waste stream production in further processes, and ﬁnally, low operating costs.
2.3.2. Magnetic Separation
Magnetic separation was selected as a possible concentration technique as the gangue minerals are
known to have lower magnetic susceptibility than value minerals which contain rare earth elements.
Based on strong or weak magnetic properties, iron-bearing minerals are characterized as ferromagnetic
or paramagnetic, respectively. Ferromagnetic refers to minerals strongly attracted to a magnet, like apiece of iron. Magnetite, maghemite, and pyrrhotite are the most common ferromagnetic minerals [20].
A laboratory-scale High Intensity Magnetic Separator (HIMS; Model 10/1, ERIEZ EUROPE,
Caerphilly, UK) with standard intensity was used to recover REE-bearing minerals (Figure 2). More than
20 tests were carried out during magnetic separation process, applying diﬀerent conditions andparameters (vibration, inclination, and speed). Two process routes were followed:
(i)
(ii)

The whole sample passing through −0.500 mm was driven to magnetic separation.
Each size fraction separately (down to +0.212 mm) was tested for magnetic separation.

The sample was fed by a vibrating feeder. The moving velocity of the feed carrying conveyor was
about 40 str/s, tilt was set initially to 72◦ . The applied magnetic intensity was 2 T. After recovering
magnetic products at 72◦ , the magnetic sample was fed again to the magnetic separator adjusted to a
74◦ and repeated to 76◦ inclination.

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Minerals 2020, 10, 387

Wet high-intensity magnetic separation (WHIMS) was also conducted on the composite A-mixedusing an ERIEZ separator by applying diﬀerent voltages, i.e., at 30 V (0.48 T) and 150 V (2.4 T).
The potential diﬀerence was selected to correspond to extreme conditions of lower and higher magnetic
ﬁeld strength in order to check the eﬀect of the intensity of the magnetic ﬁeld on the eﬃciency of themagnetic separation process. The results after magnetic separation were evaluated on the basis ofsemiquantitative mineralogical analysis using the Reference Intensity Ratio method (RIR) of I/Ic.

ab

Rich in REE (magnetic)c

Magnetic Separator HIMG ERIEZ Co. Products:
Dark color = rich in REE (magnetic)
Light color = low in REE (nonmagnetic)
Low in REE (nonmagnetic)
Figure 2. (a) Magnetic Separator HIMG ERIEZ Co. and products of separation, (b) dark coloredmagnetic product and, (c) non-magnetic product.

2.4. Evaluation of Beneﬁciation Tests
All the separated fractions from the beneﬁciation tests were tested for mineralogical and chemicalcomposition, except for the fraction +1.7 mm due to its low quantity and low content in REE.
The SEM/EDS analysis of the particle size fractions and the magnetic separates provided a roughevaluation of the test performance. The degree of liberation of the REE-bearing minerals from thegangue at various particle sizes was also established by SEM/EDS analysis. The fractions from particle
size analysis and the ﬁnal magnetic separation of the undersize 0.500 mm fractions were evaluated on
the basis of mass balance calculations.

140

Minerals 2020, 10, 387

3. Results
3.1. Mineralogy of the Sands
Table 2 and Figure S1 shows the results of mineral identiﬁcation and semiquantitative modal
composition of the composite sample “A-mixed” obtained by XRD analysis and RIR method.
The studied HM sands consist mainly of silicate minerals, Fe-oxides (magnetite, hematite), titaniteand minor amounts (<3%) of Ti oxides (ilmenite and rutile), and phosphate phases (apatite, monazite,
and xenotime) (Table 2). The major heavy minerals contained in the studied sands are amphibole(Mg-hornblende and pargasite), magnetite, titanite, allanite-epidote, and hematite (Table 2). Zircon,
monazite, cheralite, ilmenite, rutile, thorite, apatite, xenotime, baryte, and sulﬁdes were identiﬁed
as minor constituents by SEM/EDS analysis. Allanite is the major host mineral of REE in relative
abundance of 3% in the composite sample (Table 2). Based on SEM/EDS analysis and published
EPMS data, the studied monazite is classiﬁed as a cerian type (monazite-Ce). However, due tothe low abundance of the mineral as well as overlapping of monazite peak intensities with those ofmagnesiohornblende and allanite, this mineral was not clearly identiﬁed by the XRD spectra (Figure S1).
Table 2. Semiquantitative mineral contents of the composite sample (A-mixed), based on Reference
Intensity Ratio (RIR) method. Speciﬁc gravity and magnetic property data from [44–46].
MineralWeight %Nominal Speciﬁc Gravity

Magnetic Property

AlbiteQuartzK-feldspar
TitaniteMg-hornblende
Allanite
Hematite
Magnetite
Trace minerals
Total

39
31
11
8
6
3
2
1
<0.5
100

2.68
2.63
2.57
3.4–3.6
3.24
3.75
5.30
5.20
-

Diamagnetic
Diamagnetic
Diamagnetic
Paramagnetic
Paramagnetic
Paramagnetic
Ferromagnetic
Ferromagnetic-(-): not applicable.

SEM Analysis of REE Minerals(a) Allanite-(Ce)SEM/EDS semiquantitative analysis of allanite indicated that total concentration of REE oxides(TREO) range ca. from 10.7 to 16.8 wt.%, which is in agreement with the published data [18].
The stoichiometry of allanite indicates that Ce-allanite is the dominant type (Supplementary Table S2).
The backscattered election (BSE) images suggest that allanite is liberated in the particle size −0.500 mm.
The surface of the allanite is weathered with several cracks. Backscattered electron images showedthat most allanite exhibits pronounced zoning toward the marginal areas of the grain (Figure 3a,b,d).
The chemical composition of the peripheral zones indicates replacement of the allanite with epidote.
Zoned allanite contains signiﬁcant Th in the central part of the grain, ca. up to 2.5 wt.%, and they show
metamict texture. This is documented by the abundant radiation lines in the structure of the grain
displayed by this type of allanite, which results from the destructive eﬀects of its own radiation on the
crystal lattice (Figure 3).

141Minerals 2020, 10, 387

Figure 3. Backscattered electron images from SEM of various rare earth element (REE) minerals andother heavy minerals of diﬀerent particle size fractions. (a,b) zoned allanite–(Ce) showing metamict
structure and epidote peripheral zone (−0.425 mm to +0.212 mm), (c) allanite containing zircon
inclusions from the magnetic fraction, (d) zoned allanite with weathering features and cracks containing
inclusion of Ca-rich monazite (cheralite) (e) monazite containing pargasite and thorite inclusions,from the magnetic fraction, (f) thorite-weathered grain, darker zones are U depleted); abbreviations:
Aln—allanite, Ep—epidote, Zr-zircon, Mnz—monazite, Thr—thorite, Mt—magnetite, Prg—pargasite,
and Chr—cheralite.

(b) Monazite (-Ce)
Monazite occurs as very ﬁne grains, i.e., <10–20 μm in diameter, disseminated in allanite, and onlyrare liberated grains up to 200 μm are found (Figure 3b,e). In most of the cases, monazite is a
replacement phase of allanite. Monazite shows higher content in Ce than Nd or La, which agrees for acerian-type monazite and is consistent with reported data from the same area [18] (Supplementary
Table S2). ThO2 content of the monazite is generally high, i.e., 17 wt.%. The low totals reported in theanalyses are probably related to undetected MREE and HREE by SEM/EDS.
(c) Titanite
Titanite is more abundant than allanite (>5%). SEM/EDS analysis of titanites from this study
showed that the REE content is below the detection limit of the method (<0.1–0.2 wt.%). The results
are consistent with published EPMA analysis for titanites from the same area, which showed that La,
Ce, or Nd content is ≤1000 mg/kg [18].
(d) Other Heavy REE Minerals
Zircon is another REE-hosting phase, with relatively higher concentration in heavy REE (HREE)(253–890 mg/kg) relative to light REE (LREE) (8–44 mg/kg) [18,23]. Apatite is a common ﬁne-grained
accessory phase disseminated in silicate minerals. Neither apatite nor zircon or thorite was feasible
to be analyzed by SEM/EDS for REE due to the low REE content and/or the small particle size of theminerals for this method (see Supplementary Table S2).

142

Minerals 2020, 10, 387

3.2. REE Geochemistry
The chondrite-normalized REE contents of the Nea Peramos HM sand samples display similar
patterns but a large compositional range (Figure 4). The total REE (TREE) abundances are generally
higher than that of the average upper continental crust, i.e., 183 mg/kg [47], except for one
sample (#NP120) (Table 3; detailed data in Supplementary Table S1). A comparison of average
REE concentrations of the HM sands of Greece based on reported analysis [23] and HM sands
from this study shows that the Nea Peramos samples display the highest total REE enrichment.
The chondrite-normalized REE patterns of the studied samples show a pronounced enrichment oflight REE, e.g., ranging from 70 to 7000 times chondrite values, and ﬂat heavy REE (HREE) with asmall negative Eu anomaly (Figure 4). Allanite controls the LREE budget of the studied sands but theﬂat HREE characteristic is controlled by another major REE mineral, possibly titanite and/or epidote.
This characteristic is typically shown by titanite which is the major heavy mineral phase that holds
HREE (see Discussion). These observations are in agreement with previous mineralogical studies ofthe studied sands [18,23]. The average content of Sc is 7.5 mg/kg, which is lower compared to therespective value in upper continental crust, i.e., 14 mg/kg [47]. Average Y concentration in the studied
samples is 67 mg/kg and is enriched by a factor of 3 relatively to the upper continental crust, i.e.,
21 mg/kg [47]. Yttrium enrichment is associated with xenotime abundance.
Our samples present a large compositional variation in total REE (Figure 4). On the other hand,all samples show a systematic enrichment in light REE. Therefore, the preliminary beneﬁciation tests
were performed on a composite sample (A-mixed) produced by mixing of all the samples.

Figure 4. REE patterns of Nea Peramos heavy mineral (HM) sands. Red-colored line highlights themost REE-enriched sample #123. Normalization values of chondrite from [48].
Table 3. REE (Lanthanides, Sc, Y) contents of the Nea Peramos (NP) HM sands (in mg/kg). Light rareearth elements (LREE) group includes the elements from La to Sm, while heavy rare earth elements(HREE) group includes the elements from Eu to Lu; Y and Sc are included in total REE content,
according to European Union (EU) deﬁnition [2].
Sample

NP
115

NP
116

NP
117

NP
NP
117A 118

NP
119

NP
120

NP
121

NP
122

NP
123

NP
124

NP
125

NP
126

NP
127

NP
128

LREE
HREE
TREE

600
47
654

826
61
894

1407
98
1509

315
27
344

3504
204
3716

118
18
143

2024
147
2181

2225
160
2393

7645
465
8124

6646
409
7068

4552
295
4857

98
15
116

2297
76
2381

787
47
842

707
51
762

143

Minerals 2020, 10, 387

3.3. Particle Size Analysis and REE Distribution
The distribution of REE and Th in the particle size fractions of the composite sample A-mixed isdetermined by chemical assays (Table 3).
A combination of SEM/EDS and ICP-MS chemical analysis of the diﬀerent particle size fractions
showed a LREE enrichment in the fractions −0.500 to +0.212 mm and under 0.150 mm (Figure 5a).
In Table 4, mass balance calculations show that the −0.150 mm fraction, despite its composition,is insigniﬁcant in terms of beneﬁciation due to its very low mass, i.e., 1.5 wt.% Moreover, the oversize+1.70 mm particle size fraction has a very low REE content and hence is also considered as insigniﬁcant
for further beneﬁciation (Figure 5b).

Figure 5. (a) Particle size analysis and light rare earth element (LREE) distribution (mg/kg) in sample
A-mixed and (b) metal recovery in each particle size fraction, based on mass balance calculations(see text).
Table 4. Particle size analysis and REE distribution in each particle size fraction.
Particle Size
(mm)

Mass
(g)

Weight
(%)

La
Ce
Nd
Th
Concentration (mg/kg)+1.70−1.70 + 0.850−0.850 + 0.500−0.500 + 0.425−0.425 + 0.355−0.355 + 0.300−0.300 + 0.212−0.212 + 0.150−0.150
Total

55.6
230.7
640.48324.62307.32286.26353.0269.12
35.05
2302.17

2.42
10.02
27.82
14.10
13.35
12.43
15.33
3.00
1.52

16
54
416
1044
1445
1598
852
507
2031

25
106
777
1983
2719
3018
1643
968
3840

10
38
247
607
821
891
487
305
1162

7
25
190
455
641
723
390
240
945

La
0.89
12.46
266.44338.90444.08457.44300.7735.04
71.19
1927.21

Ce
Nd
Mass (mg)
1.39
24.45
497.65643.72
835.60863.93580.01
66.91
134.593648.27

0.56
8.77
158.20
197.04252.31
255.06171.9221.08
40.73
1105.66

Th
0.39
5.77
121.69147.70196.99206.97137.6816.59
33.12
866.90The maximum eﬀective recovery is achieved in the fractions −0.425 + 0.300 mm; this size range
corresponds to the “liberation” size of allanite. A similar trend of enrichment is observed in the sample#123 [17]. Thorium follows the enrichment trend of LREE and its concentration in the undersize
144

Minerals 2020, 10, 387

0.500 mm ranges from 240 to 945 mg/kg. Thorium concentration is associated with the abundance ofallanite and monazite in the HM sands, which contain 2.3 and 17–35 wt.% ThO2 , respectively, according
to EDS/SEM analysis (Supplementary Table S2).
3.4. REE Distribution in the Magnetic Fractions
High-intensity magnetic separation (HIMS) is a common separation step in REE containing beachsands in order to concentrate the targeted paramagnetic REE-bearing part and is usually applied
for monazite or xenotime [26,27]. The LREE tend to concentrate in the magnetic fraction because of
their magnetic properties. Dry HIMS was applied initially in the most REE-enriched sample #123.
Mineralogical evaluation based on XRD spectra shows that from the paramagnetic minerals allanite,titanite and magnesiohornblende tend to concentrate in the magnetic fraction under high intensities.
Qualitative evaluation of the mineralogy of magnetic and nonmagnetic separates by XRD spectra isshown in Figure 6. The peak intensities of allanite can be clearly distinguished in the magnetic separateof the sample #123 but are absent in the nonmagnetic part. Following that observation, magnetic
separation of particle size fractions from the composite A-mixed sample was carried out. The chemical
assays of the magnetic fractions are shown in Table 5 and Figure 7.
The same REE-enrichment trend, as shown in the particle size analysis, is also recorded in the HIMS
test (Figure 7). The maximum LREE concentration was achieved in the particle size: −0.355 + 0.212
and a second concentration maximum in the −0.425 + 0.355 mm fraction. Two test of HIMS wereconducted: in the ﬁrst test, all particle sizes (Feed 1 in Figure 8) were processed by magnetic separation,
whereas in the second test, only the undersize 0.500 mm were tested after removing the REE-poor+0.500 mm fraction from the initial composite sample A-mixed (preconcentration stage) (Table 5).
The results of mass balance calculations and REE distribution in every product are presented in Table 5.
Similar recoveries were achieved in the two feeds, i.e., from 92% to 87.6% and 95.9% to 78.87% for theparticle size fractions and the undersize 0.500 mm feed, respectively. The maximum LREE enrichment
is attained in the magnetic products of the diﬀerent particle sizes. It is clearly shown in the massbalance calculations in Table 5 that magnetic separation for each particle size separately improves therecovery of REE and reaches recoveries of 75–90% in just the 22% of feed material (fractions −0.500 to+0.212 mm).
Table 5. REE distribution and recovery of REE and Th in the magnetic products of each particle size of
the A-mixed sample and the undersize 0.500 mm.
Feed: Sample A (All Fractions)
Mass (g)

La

Ce

Nd

Th

La

Ce

1.87

3.29

Concentration (mg/kg)
1995

942

1647

512

365

Nd

Th

Mass (g)
1.021

0.728

Particle size

magnetic fractions (mg/kg)−1.70 + 0.850
42.55−0.850 + 0.500
162.23−0.500 + 0.425
101.65−0.425 + 0.355
109.81−0.355 + 0.300
100.6−0.300 + 0.212
142.83
Total
659.67
Recovery (wt.%)

261
1301
2764
3627
3838
3044

447
2309
4802
6321
6649
5280

137
717
1444
1865
1968
1563

93
392
1099
1472
1595
1302

0.01110.21110.28100.39830.38610.43481.722392.1

0.01900.3746
0.48810.69410.66890.75412.998991.1

0.00580.11630.14680.20480.19800.22320.895087.6

0.00400.06360.11170.16160.16050.18600.6873
94.4

Feed: −0.500 mm

La

Ce

Nd

Th

La

Ce

Nd

Th

1.92781300
95.9

3.4012
1.849389.7

Mass (g)530
Magnetic fraction
Recovery (wt.%)

Mass (g)Concentration (mg/kg)
1260
490

2223
3774

670
6089

145

460
1650

Mass (g)1.02512.983678.87

0.70380.808590.5

Minerals 2020, 10, 387

Notably, Th enrichment follows the LREE enrichment trend, and its recovery is high in theeither tested feed. The steps of magnetic separation process and the REE grade of concentrates are
summarized in the ﬂowsheet of Figure 8.

Figure 6. Comparison of XRD patterns of magnetic and nonmagnetic separates after high-intensity
magnetic separation (HIMS) from sample #123. Note that the clear peaks of allanite are absent in thenonmagnetic fraction.

Further testing on the magnetic susceptibility of the paramagnetic minerals in the composite
sample was conducted under extreme conditions of the magnetic ﬁeld, i.e., low (0.48 T) and high (2.4 T)
strength, by wet HIMS. The results are evaluated by semiquantitative analysis of the XRD spectra of
the magnetic and nonmagnetic products. The maximum content of allanite, i.e., 8 wt.%, is achieved in
the magnetic fraction at magnetic ﬁeld strength of 0.48 T (30 V) and remains constant at 2.4 T (150 V)(Supplementary Table S3).

146

Minerals 2020, 10, 387

Figure 7. (a) LREE and Th distribution in the magnetic products of the particle size fractions and(b) metal recovery in each particle size fraction, based on mass balance calculations.

Figure 8. Flowsheet of HIMS tests. Preconcentration step involves sieving and removing of +0.500 mm(Feed 1). Feed 2 corresponds to all the particle size fractions (−1.70 to +0.212 mm) of the initial composite(sample A-mixed).

147

Minerals 2020, 10, 387

4. Discussion
The beneﬁciation of REE ores is challenging from the technological point of view due to complexity
in the raw material processing [49–52]. Such processing involves three main stages: an initial
beneﬁciation step, where REE minerals are concentrated from the ore; a second step, where rare earthoxides are extracted from their host minerals producing a mixed rare earth concentrate; and ﬁnal step,
where individual rare earth metals are obtained through metallurgical separation [6].
Mineral characterization is essential for potential process development impacts. Setting thecriteria for initial target grind sizes to liberate the economic minerals and determining the possible
mineral associations among REE minerals and various gangue phases are both strongly depended on
the mineral characterization study [53]. Therefore, the information obtained herein can be used, apartfrom the prediction of the concentrate quality, in the associated concentration of deleterious elements.
X-ray diﬀraction (XRD) analysis allows the identiﬁcation of minerals. However, when mineralmodal abundance is less than ~1 wt.% considerable uncertainty and error can be introduced. It isnoteworthy that the ore grade in most REE placer deposits is commonly below 0.1% for the REE-hostmineral, e.g., monazite [54]. The evaluation of WHIMS magnetic separation products by XRD analysis
agrees with the enrichment of allanite in the ﬁnal cumulative magnetic concentrate where the content
of allanite is increased by a factor of 2.5, i.e., from 3 wt.% in the initial composite feed to ﬁnal 8 wt.%
in the magnetic concentrate (Supplementary Table S3). Correspondingly, REE contents of the ﬁnal
magnetic concentrates increased by a factor of 2.8 compared to the initial feed (Table 4, Figure 8).
This case study revealed a close association of REE-silicate minerals of allanite with monazite inclusions,as well as a high Th content in the magnetic separate associated with the REE-host minerals, e.g.,allanite-monazite/cheralite and thorite. Therefore, the information obtained herein can be used, apartfrom the prediction of the concentrate quality, in the associated concentration of deleterious elements.
One of the most serious issues associated with REE processing is the radioactive wastes produced as
by-product of the extraction stage.
The results from the present beneﬁciation tests of Nea Peramos HM sands in the course of EURARE
project also showed the high content of radionuclides, speciﬁcally Th. Thorium distribution follows theenrichment trend of LREE in the concentrates and this poses limitations in the concentration process(e.g., [10,11,23]). There are still many questions and knowledge gaps regarding the beneﬁciation
processes of the REE-bearing minerals which require a great deal of investigation [26,28]. Previous
research papers described a series of physical separation processes applied to preconcentrated rareearth minerals (REM) and discard iron oxide minerals from the magnetic fractions. These include,
gravity, magnetic, electrostatic, and ﬂotation separation techniques ([52] and references therein).
Eﬀects of Mineral Magnetic Susceptibility on the HIMS Treatment
Magnetic separation process of minerals is based on diﬀerent behavior (magnetic susceptibility)of mineral particles when exposed to an applied magnetic ﬁeld. The magnetization of a material is ameasure of the density of magnetic dipoles induced in the material [44].
The magnetic susceptibility of minerals is mainly controlled by chemical composition.
REE minerals such as allanite, monazite, bastnäsite, and xenotime exhibit moderate paramagnetic
property [44,48]. Previous studies on the beneﬁciation applicable for silicate REE minerals, including
allanite and cerite from complex ores, have shown that wet high-intensity or high-gradient magnetic
separation is more eﬀective than ﬂotation [29,55]. Yang et al. (2015) and Jordens et al. (2014) [55,56]
showed that allanite have normally poor ﬂoatability using conventional reagents.
The main host mineral for LREE in Nea Peramos HM sands is allanite-(Ce), commonly containing
inclusions of other REE-bearing phases such as thorite, zircon, monazite, and cheralite. Therefore,
REE concentration in the magnetic fraction is attributed largely to allanite which was eﬃciently
concentrated in the magnetic fraction under high-intensity magnetic ﬁeld at 0.48 T. The results ofmagnetic separation from this study are contradictory to a recent study [18]. In the previous study,

148

Minerals 2020, 10, 387

allanite was recorded to be rather separated in the nonmagnetic part but the results are not comparable
since the conditions of the applied magnetic separation test are not described by the authors.
Dry HIMS (DHIMS), applied to each particle size separately, improved signiﬁcantly the recovery of
REE with achieved recoveries of 92–88% (Table 5). Similar REE recoveries were achieved by processing
the undersize 0.500 mm, i.e., 96–79%. REE grade of the magnetic concentrate starting from the undersize
0.500 mm is better than the magnetic product of the particle size fractions. In general, the actual TREE
concentration in the magnetic products was 2.8 times higher compared to the initial feed contents, inthe two feeds tested. The optimum magnetic separation was eﬀected after repetitive passages of themagnetic fraction at increasing belt inclination from 72◦ to 76◦ . Moreover, the preliminary results of
WHIMS corroborated for an eﬀective recovery of allanite in the magnetic fraction, and the enrichment
achieved was in the order of 2.5 times in the magnetic product (Supplementary Table S3).
Additional laboratory tests for REE dissolution and extraction from concentrate was tried by acidtreatment, either leaching or acid baking followed by water leaching conducted by I.G.M.E. during
EURARE project [57]. The acids used were HCl, H2 SO4 , or both. The results showed that rather low
recoveries were achieved. In direct acid leaching, decreasing pulp density leads to increasing rare earth
recovery, while the opposite is observed in the recovery procedure with acid baking; an increasing pulp
density at acid baking step at about 15% leads to an increasing recovery tendency of about 15–20%(unpublished data, Angelatou pers. comm.). The duration of acid baking test seems to have no eﬀect
on the recovery of rare earth elements.
5. Conclusions
HM sands from the shoreline of Nea Peramos, Kavala, Northern Greece was studied in order to beable to determine the best process to concentrate the ore by testing simple screening as a prebeneﬁcation
method and high-intensity magnetic separation as the main beneﬁciation method. The results of thetest work lead to the following conclusions:
•

•
•
•
•
•
•
•

Allanite-(Ce) is the major host mineral for light REE (LREE), whereas monazite, zircon, and thorite
constitute trace amounts. Titanite displays low concentration in LREE <0.1–0.2%. Metamict
allanite is common and is thorium-enriched relative to the nonmetamict allanite.
A simple screening can achieve a satisfactory prebeneﬁcation.
A stepwise magnetic separation improves the recovery of REE.
Magnetic separation for each particle size fraction separately improves the recovery of REE andreaches recoveries of 75–90% in just the 20% of feed material.
The grades of magnetic concentrate for the two processes (all particle size fractions and undersize
0.500 mm) were 0.84% and 1.15% TREE, respectively, at the recoveries of 87–92% and 79–96%.
The increase in REE content is associated with the increase of thorium content in concentrates
regarding placer sands from Nea Peramos, Greece.
The use of gravimetric methods (such as Wilﬂey shaking table) did not contribute much to the
beneﬁciation of the ores.
Radioactive wastes from a potential REE processing operation of Nea Peramos HM sands should
be carefully encountered for the risk to the environment and humans.

Supplementary Materials: The following are available online at http://www.mdpi.com/2075-163X/10/5/387/s1,
Table S1: REE contents of the Nea Peramos HM sands (in mg/kg), Table S2: SEM/EDS spot analysis of REE-bearing
minerals, and Table S3: Semiquantitative mineralogical composition of the wet HIMS-treated samples and thecomposition of the initial A-mixed composite sample for comparison. Applied voltage values of 30 and 150 V
correspond to 0.48 and 150 T magnetic ﬁeld strength, respectively, Figure S1: X-ray diﬀraction patterns of composite
A-mixed and sample #123, in the range of 20–43◦ showing in magniﬁcation the identiﬁed heavy minerals. Quartz
appears as the stronger peak. The major peaks of intensities for the heavy minerals only are marked. Monazite(-Ce) peaks shown are indicative (database for RRUFF™ Project [43]). Ttn = titanite, Mz = monazite, Mg-Hbl =magnesiohornblende, Prg = pargasite, Aln = allanite, Mt = magnetite, Hem = hematite).

149

Minerals 2020, 10, 387

Author Contributions: V.A. and D.E. collected the samples and provided the ﬁeld information; V.A. conceived
and designed the experiments undertaken by I.G.M.E. in the framework of EURARE project; V.A. and S.P.
performed the experiments; C.S. supervised the mineralogical analysis; and C.S., S.P., V.A. and K.S. wrote thepaper. All authors have read and agreed to the published version of the manuscript.
Funding: Part of this research was supported by EURARE research project, funded from the European Community’s
Seventh Framework Programme ([FP7/2007–2013]) under grant agreement no 309373. Project web site: www.
eurare.eu.
Acknowledgments: The authors would like to thank N. Xirokostas for elaborating the chemical analysis of REE
ore at the ICP-MS facility of I.G.M.E., V. Skounakis for assistance in performing the SEM/EDS analysis at NKUA,
Geology and Geoenvironment, Athens, and I. Marantos from I.G.M.E. for assistance in the quantitative evaluationof the XRD patterns by RIR method. This publication reﬂects only the authors’ view, exempting the Communityfrom any liability.
Conﬂicts of Interest: The authors declare no conﬂict of interest.

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© 2020 by the authors. Licensee MDPI, Basel, Switzerland. This article is an open accessarticle distributed under the terms and conditions of the Creative Commons Attribution(CC BY) license (http://creativecommons.org/licenses/by/4.0/).

152minerals
ArticleFactors Controlling the Chromium Isotope
Compositions in Podiform Chromitites
Maria Economou-Eliopoulos 1, *, Robert Frei 2 and Ioannis Mitsis 1
1
2*

Department of Geology and Geoenvironment, University of Athens, 15784 Athens, Greece;
mitsis@geol.uoa.gr
Department of Geosciences and Natural Resource Management, University of Copenhagen, Øster Voldgade
10, 1350 Copenhagen K, Denmark; robertf@ign.ku.dk
Correspondence: econom@geol.uoa.gr

Received: 12 November 2019; Accepted: 19 December 2019; Published: 21 December 2019

Abstract: The application of Cr isotope compositions to the investigation of magmatic andpost-magmatic eﬀects on chromitites is unexplored. This study presents and compiles the ﬁrst Cr
stable isotope data (δ53 Cr values) with major and trace element, contents from the Balkan Peninsula,aiming to provide an overview of the compositional variations of δ53 Cr values in ophiolite-hosted
chromitites and to delineate geochemical constraints controlling the composition of chromitites. The
studied chromitites exhibit δ53 Cr values ranging from −0.184% to +0.159%, falling in the range ofso-called “igneous Earth” or “Earth’s mantle inventory” with values −0.12 ± 0.11% to 0.079 ± 0.129%
(2sd). A characteristic feature is the slightly positively fractionated δ53 Cr values of all chromitite
samples from Othrys (+0.043 ± 0.03%), and the occurrence of a wide range of δ53 Cr values spanningfrom positively, slightly negatively to the most negatively fractionated signatures (Pindos, δ53 Cr =−0.147 to +0.009%; Skyros, δ53 Cr = −0.078 to +0.159%). The observed negative trend between δ53 Cr
values and Cr/(Cr + Al) ratios may reﬂect a decrease in the δ53 Cr values of chromitites with increasing
partial melting degree. Alternatively, it may point to processes related to magmatic diﬀerentiation, as
can be seen in our data from Mikrokleisoura (Vourinos).
Keywords: chromium isotopes; chromitites; ophiolites; Balkan Pensula

1. Introduction
Chromitites in ophiolite complexes are studied extensively, in particular with respect tomantle-lithospheric slab interactions, post magmatic processes and subduction recycling scenarios [1–8].
Chromitites diﬀer in their major and trace element compositions, including their platinum-groupelement (PGE) content, and also have been shown to exhibit variations in their chromium isotope
signatures. Although the PGE content in chromitites is commonly a few 100 s of ppb, a signiﬁcant
PGE-enrichment, reaching tens of ppm has been reported in high-Cr or high-Al type chromitites
from small occurrences, either in Pt and Pd (PPGE; incompatible, with Di > 1) or in Os, Ir, Ru (IPGE;
compatible, Di < 1). Primary magmatic compositional trends recorded in chromite may be obliterated
later on during secondary processes, in that Platinum-Group Minenerals (PGM) and silicates may
have been substantially modiﬁed during subsequent subsolidus re-equilibration and/or by reaction
with metasomatic ﬂuids, during extended periods of deformation (including ductile asthenospheremantle ﬂow), and by shallow crustal, brittle deformation. The factors controlling the trace element
distribution in chromitites remain unclear [8–20].
Likewise, limited data on the chromium isotope compositions of chromitites have shown
isotopically heavier signatures compared to mantle peridotites [2,21–23]. However, the role ofmagmatic processes on the Cr isotope system during partial melting and magma diﬀerentiation isMinerals 2020, 10, 10; doi:10.3390/min10010010

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still unexplored, and there is still extensive debate on the reasons for the diversity of PGE contents inchromitites [22–24].
Although δ53 Cr values of chromitites from the Balkan peninsula are still very limited [2], Cr isotope
studies have been performed by several authors in the view to evaluate hexavalent chromium (Cr(VI))
contamination in groundwater and rock leachates [25–30]. For example, Cr isotope signatures indicatedthat Cr(VI) contamination in groundwater from the Friuli Venezia Giulia Region of Italy is related tothe oxidation of trivalent (Cr(III)) deposited as a consequence of past industrial activities [25]. The
δ53 Cr values and Cr(VI) concentrations in contaminated water from the Evia and Assopos Basins(Greece) are compatible with global data ranges that characterize waters contaminated by both naturalprocesses and anthropogenic activities, and probably delineate potential contamination sources [28].
Also, Cr isotope-based studies have shown oxidative mobilization of Cr(VI) from ultramaﬁc host rocks,
and successive back-reduction of the thus mobilized Cr(VI) fractions [29]. Furthermore, the record of a
12-month long time-series of δ53 Cr values in run-oﬀ from a small serpentinite-dominated catchment
in Central Europe, revealed that the Cr(VI) export ﬂux during winter was signiﬁcantly higher thanduring the summer [29]. These recent studies and respectively supporting data, in the light of thewidespread distribution of ophiolite complexes in Europe and in orogenic zones elsewhere around theglobe, emphasize the need for further more detailed studies addressing the variation of δ53 Cr values in
Cr-bearing rocks and ores.
The ophiolite complexes form an important component of the Tethyan metallogenetic belt,
extending from the Mirdita zone in Albania to the Vourinos, Pindos and Othrys complexes ofthe south zone. These ophiolite complexes host large chromite deposits and smaller chromite
occurrences, characterized by high-Cr and high-Al, low- and high-IPGE and PPGE compositions, andby PGE-patterns with both negative and positive slopes [10,14,15,31]. Such a compositional diversity
in chromitites may provide an opportunity for in detail studies aimed at a better understandingof the role of magmatic processes on the Cr isotope signatures of the respective chromites [21–24].
In our study we present the ﬁrst systematic chromium isotope data (expressed as δ53 Cr values) of
massive chromitites hosted in ophiolite complexes of the Balkan Peninsula, along with trace element
contents, including PGE, and combined with scanning electron microscopy/energy-dispersive X-ray
spectroscopy (SEM/EDS) identiﬁcations. We aim at delineating geochemical constraints that potentiallycontrol the elemental composition of these chromitites and their Cr isotope compositions, ultimately
with the overall intention to contribute to the origin and genesis of chromite and PGE mineralizationand deposits formation.
2. Materials and Methods
2.1. Preparation and Analysis of the Chromitites
For the purpose of the present study, 2.5 to 3.0 kg amounts of chromitite sample mostly fromlocations of the Balkan Peninsula (Table 1) were crushed into pieces of approximately 0.3 × 0.3 cm withthe help of a jaw crusher. The crushed piles were mixed well and divided into four equal portions.
Subsequently, two fractions were removed and the remains were remixed again. The procedure wasrepeated until the ﬁnal split weighed about 500 g. This split, serving as experimental material, waspulverized to about 100 mesh using an agate mortar. A conventional oxidative alkaline fusion (OAF)
was carried out in corundum crucibles; about 1 g of sodium peroxide (Na2 O2) and 0.3 g of Na2 CO3
were added to about 0.1 to 0.3 g of the powdered sample (amount depended on the concentration of Cr
in the sample). The sample and ﬂux materials were mixed and molten at temperatures of 700 to 800 ◦ C
in a muﬄe furnace during 10 min. After cooling, the fused cake was extracted from the crucible andtransferred into 100 mL volumetric ﬂasks using deionized water [2]. The solutions were ﬁnally ﬁlteredthrough 0.45 μm polyamide membrane ﬁlters and aliquots of these solutions were then processed forchromium isotope analysis.

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2.2. Chromium Isotope Analysis
Solutions of samples in the amount which would yield about 1 μg of total chromium were pipetted
into 13 mL Savillex Teﬂon beakers. These aliquots were spiked up with adequate amount of a 50 Cr–54 Cr
double spike so that a sample to spike ratio of ~3:1 (total chromium concentrations) was achieved.
The addition of a 50 Cr–54 Cr double spike of a known isotope composition to a sample before chemical
puriﬁcation allows accurate correction of both the chemical and the instrumental shifts in Cr isotope
abundances [1,32]. The mixture was totally evaporated and 1 mL of concentrated aqua regia wassubsequently added. After 3 h during which the sample was exposed to aqua regia on a hotplate at
100 ◦ C, the sample was again dried down. The sample was then puriﬁed by passing the sample in
0.5 N HCl over an extraction column (BioRad PP columns) charged with 1 mL of 200–400 mesh BioRad
AG-50W-X12 cation resin, employing a slightly modiﬁed extraction recipe published [33,34]. The Cr
yield of this column extraction and puriﬁcation step is usually ~70%. Samples were loaded onto Re
ﬁlaments with a mixture of 3 μL silica gel, 0.3 μL 0.5 mol L−1 of H3 BO3 and 0.5 μL 0.5 mol L−1 of H3 PO4.
The samples were statically measured on a IsotopX “Phoenix” multicollector thermal ionization mass
spectrometer (TIMS) at the Department of Geoscience and Natural Resource Management, Universityof Copenhagen, at temperatures between 1040 and 1150 ◦ C, aiming for beam intensity at atomic massunit (AMU) 53.9407 of 1 to 2 Volts. Every load was analyzed ﬁve times. Titanium, vanadium and ironinterferences with Cr isotopes were corrected by comparing with 49 Ti/50 Ti, 50 V/51 V and 54 Fe/56 Fe ratios.
The ﬁnal isotope composition of a sample was determined as the average of the repeated analyses andreported relative to the certiﬁed (standard reference material) SRM 979 standard as
δ53 Cr(%) = [(53 Cr/52 Crsample /53 Cr/52 CrSRM979 ) − 1] × 1000
Repeated analysis of 1 μg loads of unprocessed double spiked SRM 979 standard during theduration of the analysis period yielded an average δ53 Cr value of 0.05 ± 0.06% (n = 12; 2σ; 52 Cr signal
intensity at 2 V) on the “Phoenix” TIMS which we consider as a minimum external reproducibility for asample run in this study, including separation procedure, double spike correction error, and respective
internal analytical errors.
2.3. Mineral Chemistry and Whole Rock Analyses
Polished sections from chromite deposits of Greece (Table 1) were carbon or gold coated andexamined by reﬂected light microscopy and with a SEM using EDS. SEM images and EDS analyses
were carried out at the University of Athens (Department of Geology and Geoenvironment) using
a JEOL JSM 5600 instrument, equipped with automated energy-dispersive analysis system ISIS 300
OXFORD, with the following operating conditions: accelerating voltage 20 kV, beam current 0.5 nA,time of measurement (dead time) 50 s and beam diameter 1–2 μm. The following X-ray lines wereused: OsMα, PtMα, IrMα, AuMα, AgLα, AsLα, FeKα, NiKα, CoKα, CuKα, CrKα, AlKα, TiKα, CaKα,SiKα, MnKα, MgKα, ClKα. Standards used were pure metals for the elements Cr, Fe, Mn, Ni, Co, Ti
and Si, MgO for Mg and Al2 O3 for Al. Contents of Fe2 O3 and FeO were calculated on the basis of thespinel stoichiometry.
Major and trace elements in massive chromitite samples were determined by inductively coupledplasma mass spectrometry (ICP-MS) analysis, at the ACME Analytical Laboratories Ltd., Vancouver,
BC, Canada. The samples were dissolved using a strong multi-acid (HNO3 –HClO4 –HF) digestion andthe residues dissolved in concentrated HCl. Platinum-group element (PGE) analyses were carried out
by Ni-sulphide ﬁre-assay pre-concentration technique, using the nickel ﬁre assay technique from large(30 g) samples. This method allows for complete dissolution of samples. Detection limits are 5 ppb for
Ru 2 ppb for Os, Ir, Pt, Pd and 1 ppb for Rh and Au. The CDN-PGMS-23 was used as standard.

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3. A Brief Outline of Characteristics of Ophiolites and Hosted Chromitites
3.1. Ophiolites
The geology, petrography, mineral chemistry and geochemistry of ophiolites in the Balkan
Peninsula has been a topic of extensive investigation in previous publications [10,35–48]. They are
an important component of the Upper Jurassic to Lower Cretaceous Tethyan ophiolite belt, which
extends through the Serbian zone of the Dinarides in the north (Mirdita zone) to the Subpelagonianzone (Pindos, Vourinos and Othrys complexes) in the south (Figure 1).

Figure 1. Simpliﬁed map of the geotectonic zones of Hellenides [35], showing the ophiolite complexes
from which chromitite samples were studied. Symbols: Oph = ophiolites; Rh = Rhodope zone; Sm =Serbo–Macedonian Massif; CR = Circun–Rhodope zone; (Pe, Pa, Al) = Axios zone; Pl = Pelagonian
zone; AC = Attico–Cucladic zone; Sp = Pindos zone; P = Parnassos–Giona zone; G = Gavrovou–Tripolis
zone; I = Ionian zone; Px = Pakson zone.

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These ophiolites are characterized by petrological and geochemical features typical of both
Mid-Ocean Ridge (MOR) and Supra-Subduction Zone (SSZ)-type associations (fore-arc and back-arc
ridges). They are accompanied by minor dunite bodies, and overlain by gabbroic cumulates, maﬁc dykes,
and an extrusive sequence with a compositional basalt range between MOR and island-arc-tholeiitic toboninitic types [35–48]. Recently, ophiolite studies have focused on the importance of intra-oceanic
subduction-initiation processes in ophiolite genesis [49]. Relatively small isolated ophiolite masses,of mostly serpentinized dunite and harzburgite, are located at the western margin of the Axios
zone (Vermio–Veria), and the Eohellenic Pre-Cretaceous nape, including Skyros Island (Figure 1).
Structural and paleomagnetic studies on those ophiolites have revealed widespread heterogeneous
deformation and rotation during their original displacement and subsequent tectonic incorporation into
continental margins [35–38]. The origin of major Jurassic ophiolite complexes in Greece in a hydrous
SSZ environment has been evidenced by the presence of hydrous silicate inclusions, such as amphibole
and phlogopite, in chromitites and by the trace and Rare Earth Elements (REE) data on separated
orthopyroxenes and clinopyroxenes from harzburgites of these ophiolites [48]. Pyroxenites withvariable modal contents of olivine, garnet and spinel, ranging in composition from orthopyroxenite
through websterite to clinopyroxenite have been described in the Veria ophiolites [50,51].
3.2. Characteristics of Chromitites
Chromitites in the Balkan Peninsula are classiﬁed on the basis of the variation in the chromitite
tonnage, the composition of chromite, the degree of transformation of ores and the associated ophiolites,
and on the association of chromitites with sulphides.
The Othrys complex includes two tectonically separated chromite deposits, namely Eretria (Tsagli)
and Domokos, in addition to several other occurrences, including Agios Stefanos, the combined
tonnage being approximately 3 Mt of high-Al massive chromite ores occurring in moderately depleted
harzburgite [10,35,39,44,45]. A salient feature of the Eretria (Tsagli) chromite deposits is the occurrenceof Fe–Ni–Cu-sulphide mineralization with dominant minerals pyrrhotite, chalcopyrite and minor
cobalt-bearing pentlandite, hosted in serpentinized harzburgites [52]. Massive sulphide mineralization
occurs at the peripheral parts of podiform chromite bodies in association with chromitite and magnetite(Figure 2). Chromite is often characterized by brittle deformation and by sulphide-silicate veins(Figure 2d,e). Texture and geochemical characteristics, including PGE content, ﬂat chondrite-normalized
PGE-patterns, and very low partition coeﬃcients for Ni and Fe between olivine and sulphides
are inconsistent with sulphides having been in equilibrium with Ni-rich host rocks, at magmatic
temperature [52,53]. In addition, it has been documented [54] that in the Othrys and Vourinos
ophiolite complexes massive and disseminated chromitites host considerable amounts of methane in
micro-fractures and in porous serpentine- or chlorite-ﬁlled veins.

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Figure 2. Photographs showing the texture of chromitite associated with sulphides and magnetitefrom the peripheral parts of podiform chromitite bodies at Eretria. Chromite occurs interstitially insulphide rich portions (a,b) and sulphides occur as veins and irregular masses within zones dominatedby chromite (b,c). Back-scattered electron images from massive chromitite samples from central partsof podiform bodies, showing fragmented chromite and evidence for the existence of sulphides (d,e)
even distant from the peripheral parts of the chromitite bodies. Abbreviations: chr = chromite; sf =sulphides; po = pyhhrotite; srp = serpentine; chl = chlorite; grt = garnet. Scale bar: 1 cm is the same forall hand samples.

Chromite ores in the Vourinos complex occur in tectonite and cumulate sequences. Only those in
the tectonites are under exploitation. All chromite ores, with a tonnage estimated to be approximately
10 Mt of high-Cr ore are found within dunite bodies or dunite envelopes in harzburgite, but there is no
systematic relation between size of the dunite body and that of the ore body. The sizes of the ore bodies
vary widely and contain all textural types (massive, schlieren, banded, disseminated and nodular), but
usually a single type dominates. High-temperature deformation, superimposed on primary magmatic
textures is very common [42,47]. The mantle sequence of the Pindos ophiolite complex resembles thatof Vourinos in the presence of extensive and highly depleted harzburgite but, in contrast to Vourinos,
there is only a limited number of small chromite occurrences (low potential for exploitation).
The chromitite occurrences in the Pindos ophiolite complex are small (a few tens of meters × a
few tens of centimeters) and cover all textural types (massive, schlieren, banded, disseminated andnodular). Chromitites occur within completely serpentinized and weathered, intensively deformed
dunite–harzburgite blocks, as a result of strong plastic and brittle deformation that was superimposedon primary magmatic textures. The chromitites are mostly ﬁne- to medium-grained, and consist of
aggregates of fractured chromite accompanied by chlorite and traces of tremolite. Primary olivine ispreserved only in the form of inclusions within chromian spinel, while abundant remnants of basemetal sulphides (BMS), now preserved as alloys, occur both as inclusions and interstitial phases inthe silicate matrix [10,18,19]. Chromitites throughout the Pindos complex are of high-Cr and high-Al
compositions, often in a spatial association with each other [45,55–57]. The most salient feature of thePindos chromitites is the PPGE-enrichment in those bodies occurring in the area of Korydallos, with
PGEtotal concentrations reaching 7 ppm [43–45] and reportedly up to 29 ppm [18].
Relatively small massive chromitite bodies are located at the Achladones area on the island of
Skyros. They are of high-Al type (average Cr/(Cr + Al) ratio of unaltered chromite is 0.56 while theMg/(Mg + Fe2+ ) ratio is 0.64) and contain elevated PGE contents, up to 3 ppm PGEtotal . However,158

Minerals 2020, 10, 10

both high-Cr and high-Al types have been reported from Skyros [46,47,58]. Furthermore, SEM/EDS
investigations of Au-coated polished sections of chromitite from Skyros revealed the presence ofgraphite (Figure 3).
Ophiolites associated with the Serbomacedonian massif (Gomati and Exochi) and Rhodope
massif including the ophiolites Souﬂi–Tsoutoura (Greece), Dobromirci, Jacovitsa, Broucevci and
Goliamo–Kamenjane (Bulgaria) (Figure 1) are completely serpentinized, locally sheared andmetamorphosed to antigorite–tremolite and/or talc schists, and they are all small (a few thousand tonsof ore). A characteristic feature of those chromitite occurrences is the common spatial associationof high-Cr and high-Al ores, and the negative slope of the PGE patterns, comparable to those for
podiform chromitites in ophiolites elsewhere [59–61]. The area of Ceruja is part of the western Bulqiza
ophiolite complex as part of the Mirdita zone (Albanides), and includes harzburgitic to lherzolitic
types [62]. Although more than 40 Mt of chromitites have been located in the uppermost part of the
mantle harzburgites throughout the Bulqiza complex, in the area of Ceruja there are only small (a few
thousand) chromite occurrences.
There is a wide compositional variation in chromite throughout the Bulqiza complex, with Cr/(Cr+ Al) ranging from 0.46 to 0.86 and Mg/(Mg + Fe2+ ) ratios ranging from 0.46 to 0.69. The Cr/Cr +Al) ratio in studied samples from the area of Ceruja is restricted to 0.43–0.56, their PGE content is
low and the PGE-patterns show a negative slope, similar to other chromitites hosted in ophiolitecomplexes [62]. However, in the western Bulqiza massif, mineralization occurring as disseminations
in dunite is accompanied by chromite in the transition zone and lower cumulates sequence of thecomplex and elevated PGE contents, up to 9000 ppb ΣPGE [62–64].

Figure 3. Back-scattered electron images (Au-coated polished sections) of strongly fragmented
chromitite from Skyros Island, showing microstructural inclusions of graphite. Partial replacement ofserpentine ﬂakes by graphite (a,b,d); Overgrowth of ﬁne graphite nodules (white arrows) are revealedon unpolished parts of polished sections (a,c); (f) Dissolution pits in silicate inclusions of chromite are
ﬁlled with graphite (b,d). Abbreviations: grp = graphite-like chr = chromite; srp = serpentine; chl= chlorite.

4. Results
4.1. Distribution of Selected Trace Elements
Selected trace element contens of Ni, Co, Mn, Zn, V, Ga Ti, and platinum-group elements, which
are from previous studies, are contained in Table 1.

1590.023
0.036
0.025
0.088

−0.12
−0.093−0.078−0.056−0.032−0.03
0.009−0.096−0.149−0.0980.062−0.102−0.06−0.086−0.147−0.0780.117
0.159−0.111−0.084−0.051−0.129−0.047

Othrys (D), GR
Eretria
Eretria
Domokos
Domokos

Vourinos (D), GR
KondroVoidolakos
Kissavos
Kissavos
Mikrokleisoura
Mikrokleisoura

Pindos, Low PGE
P Dako’s mine
Korydallos
Kampos Despoti
Korydallos

Pindos, High-PPGE
Korydallos
Korydallos
Korydallos
Korydallos
Korydallos

Pindos, High-IPGE
Milia

Skyros, high IPGE
Achladones
Achladones
Achladones

Skyros, Low PGE
Ag. Ioannis
Parapisti
CerujaCerujaCeruja

1
2
3
4

5
6
7
8
9
10

11
12
13
14

15
16
17
18
19

20

160

21
22
23

24
25
26
27
28

0.75
0.79
0.43
0.52
0.56

0.61
0.55
0.58

0.82

0.68
0.67
0.69
0.4
0.4

0.52
0.48
0.81
0.42

0.77
0.81
0.81
0.56
0.78
0.61

0.57
0.56
0.53
0.5

0.69
0.69
0.69
0.58
0.65

0.64
0.69
0.51

0.42

0.57
0.57
0.49
0.72
0.72

0.63
0.6
0.58
0.71

0.65
0.62
0.66
0.72
0.61
0.52

0.63
0.66
0.74
0.73

Mg

SEM/EDSCr(%)

LocationNo

δ53 Cr

30
25
3
3
3

140
60
13

150

14
47
62
88
45

4
40
30
8

26
14
14
4
30
50

22
25
15
20

Os

15
20
7
7
7

480
40
31

320

11
49
47
74
38

6
20
18
3

36
17
11
3
11
8

13
14
11
10

Ir

85
90
18
15
18

1200
300
159

350

57
55
80
130
82

30
33
50
11

55
80
49
11
40
42

67
25
30
35

Ru

6
5
3
2
3

160
13
7

82

13
104
112
142
78

1
5
3
1

17
12
7
3
16
18

2
3
5
5

Rh

PGE Content (ppb)

4
3
25
28
25

280
23
20

150

3460
3020
1460
1720
4580

4
25
76
75

4
6
4
3
8
8

10
4
24
8

Pt

5
2
7
9
7

39
28
21

7

1660
600
337
525
1300

6
20
4
19

2
6
7
6
4
6

1
1
2
11

Pd

145
145
63
64
63

2300
464
251

1059

5215
3875
2098
2679
6123

51
143
181
117

140
135
92
30
109
132

115
72
87
89

0.33
0.1
1
1.3
1

0.08
0.7
0.67

0.02

151
12.2
7.2
7.1
34.2

1.0
1.0
0.22
6.33

0.06
0.35
0.55
0.54
0.63
0.75

0.08
0.07
0.18
1.1

ΣPGE Pd/Ir

0.23
0.3
1.74
2.1
1.74

1.13
0.39
0.53

2.00

7.53
3.86
2.4
2.01
4.6

0.52
0.8
7.01
5.5

0.22
0.23
0.55
0.23
0.26
0.96

2.26
0.74
2.43
0.34

Pt/Pt *

1250
1200
2640
1640
2150

1300
1600
1500

1600

1550
750
720
630
1630

1300
1500
1450
1500

2000
1580
1900
1800
1800
1000

1600
1400
1600
1200

Ni

220
200
170
160
160

250
200
240

340

230
270
240
240
240

140
260
290
260

240
210
200
170
180
230

640
620
610
1030
990

1200
1000
870

980

510
760
760
790
580

600
710
560
710

500
560
400
780
620
1000

900
1100
960
900

V

420
400
300
420
340

540
400
450

1160

450
520
620
510
460

400
410
490
520

260
550
300
360
280
490

340
430
310
400

Zn

2150
1900
960
1210
1100

1300
1650
1600

2000

1200
1280
1340
1300
1400

760
1200
1470
1300

1560
1540
1000
850
950
1760

1250
990
1300
2500

Mn

14
11

34
36
40

15

30
16
23
25
32

23
32
17
34

14
14
13
27
14
20

34
32
32
33

Ga

Trace Element Content (ppm)

220
210
200
220

Co

Table 1. Chromium isotope (δ53 Cr) and selected trace element data for chromitites from the Balkan Peninsula.

6
5

5
5
5

5

5.7
6.8
6.7
6.8
5.4

5
5
5
5.1

6
6
6
6
8
12

6
6
7
6

Sc

11
10

11
9.4
9.6

12

7.8
9.7
10
9.9
7.3

7.8
9.8
10
7.8

8.1
7.9
8.7
9.2
7.8
15

9.6
10
9.8
9.6

Fe

wt %

<0.1
<0.1

<0.1
<0.1
<0.1

0.1

<0.1
<0.1
<0.1
<0.1
<0.1

<0.1
<0.1
<0.1
<0.1

<0.1
<0.1
<0.1
<0.1
<0.1
0.5

<0.1
0.2
<0.1
<0.1

Ca

wt %

0.08
0.05

0.04
0.04
0.04

0.01

0.07
0.07
0.08
0.09
0.07

0.04
0.08
0.03
0.09

0.04
0.04
0.04
0.03
0.05
0.1

0.03
0.03
0.04
0.04

Ti

wt %

Minerals 2020, 10, 10

38
39
40

36
37

35

32
33
34

0.6

0.02
−0.0570.084

SMM, GR
Exochi
Exochi
Gomati

0.62
0.6
0.67

0.72
0.61

0.62

0.62
0.65
0.68

0.66

0.62
0.65

23
25
9

5
22

3

20
25
16

70

70
30

Os

24
10
20

13
42

6

25
6
32

15

15
20

Ir

73
70
60

23
53

14

85
40
120

100

100
55

Ru

6
4
6

4
7

3

7
2
6

9

9
4

Rh

PGE Content (ppb)

5
4
5

2
3

3

7
3
4

3

3
4

Pt

7
6
4

12
2

1

5
6
14

2

2
31

Pd

140
119
104

60
129

30

150
82
190

199

199
144

0.29
0.6
0.25

0.92
0.05

0.2
1.00
0.44

0.13

0.13
0.05

ΣPGE Pd/Ir

Table 1. Cont.

0.25
0.26
0.33

0.09
0.26

0.38
0.28
0.14

0.23

0.23
0.12

Pt/Pt *

1240
1030
1050

800
790

2000

1260
1430
520

1770

1500
880

Ni

240
150
330

180
240

780

210
220
160

150

670
1300
660

600
650

500

610
670
590

550

650
700

V

400
300
350

370
400

3700

360
360
230

300

700
500

Zn

1850
1270
2700

2500
3500

7500

1850
1270
2300

1800

2800
1400

Mn

Ga

Trace Element Content (ppm)

200
170

Co

* Present data; PGE data from previous publications [24,34,43,44,50]. Symbols: Cr# = Cr/(Cr + Al); Mg# = (Mg + Fe2+ ).

0.76
0.75
0.49

0.7
0.63
0.65

−0.037−0.033−0.184

0.47
0.75

0.75

−0.145−0.0990.75
0.7−0.075−0.076−0.147−0.09729
30

Mg

SEM/EDSCr(%)

Rhodope massif, Bu
Souﬂi
Tsoutoura
Dobromirci
GoliamoKamenjane
Broucevci
Jacovitsa

Vermio-Veria
Veria
VermioRhodope massif,
GR

31

LocationNo

δ53 Cr
Sc

Fe

wt %
Ca

wt %

Ti

wt %

Minerals 2020, 10, 10161

Minerals 2020, 10, 10

The rare earth elements La, Ce, Nd, Sm, Yb, Pr, Eu, Gd, Tb, Dy, Ho, Er, Tm and Lu as well as
U, Th, Cs, Sb, Y, Zr, Ge, W, Mo, Sc, Pb, Li, Sr and K were lower than the detection limits in detailspeciﬁed in the section of the analytical methods, except for La and Ce contents in the chromitite
samples from Othrys and Skyros island, which are slightly higher than the respective detection limit.
The best deﬁned inverse correlations are those between the Cr/(Cr + Al) ratio and V (r = –0.80 to–0.85) for chromitites from Vourinos, Pindos (low PGE) and Skyros, and between the Cr/(Cr + Al) ratio(Figure 4a), and Ga (r = –0.90 to –0.96) for chromitites from Vourinos, Pindos (low PGE) and Skyros
and for high-PGE chromitites (r = –0.65; Figure 4b).

Figure 4. Plots of the Cr/(Cr + Al) ratio versus V and Ga for chromites from the Balkan Peninsula. The
best pronounced negative correlations are those between Cr/(Cr + Al) ratio and V for chromitites from
Vourinos, Pindos (low PGE) and Skyros (a), and Ga for chromitites from Vourinos, Pindos (low PGE)
and Skyros (b). Data from Table 1.

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4.2. Chromium Isotope Results
The analysed chromitites, representative of a wide range in their major element (high-Cr andhigh-Al) and PGE composition, show a relatively wide range of chromium isotope compositions,expressed as δ53 Cr values, even in spatially related occurrences, throughout the entire metallogenic belt.
The δ53 Cr values range from −0.184% (in the Bulgarian Rhodope massif) to +0.159% on Skyros Island(Figure 5; Table 1). The lower values (−0.246%) are comparable to those having been reported by [2].
δ53 Cr as low as −0.21%o are compatible with the δ53 Cr isotope range for mantle-derived rocks [1],while the more positively fractionated values of up to +0.159%o exceed those values around +0.068%
reported for global chromitites from layered intrusions and podiform-type occurrences [2].

Figure 5. δ53 Cr values of the chromitites (n = 40) from the Balkan Peninsula (Table 1) studied herein.
The blue coloured band displays the range of high-temperature magmatic signatures as has been
deﬁned by [1]. Errors are 2sd. The reproducibility of the (standard reference material) SRM 979
standard under similar measuring conditions is at +/−0.06%. The horizontal axis corresponds to thenumber of analyzed samples 1–40 (Table 1).

Along with the δ53 Cr values of the chromitites, selected major and trace element compositions
reported in previous publications are given (Table 1) in order to delineate potential geochemical
constraints that control their composition and the Cr isotope compositions. It is remarkable that
all chromite samples from Othrys and certain ores from Skyros exhibit positive δ53 Cr values, while
chromitites from the Serbo-macedonian-Rhodope massifs and Pindos show a wide range from positive,
slightly negative to the most negative δ53 Cr values (Figures 5 and 6; Table 1). Also, the results do
not show any correlation trend either between total PGE content or Mg/(Mg + Fe2+ ) ratio and δ53 Cr
values. There is a slightly negative trend between δ53 Cr values and Cr/(Cr + Al) ratio, which is better
pronounced for the chromitite samples from Othrys and Vourinos (Figure 6a) and a positive trend
between δ53 Cr values and Ga content (Figure 6b).
The Pd/Ir values for chromitites from the Balkan Peninsula are plotted versus normalized values(Pt/Pt* = (Pt/8.3)/(Rh/1.6)(Pd/4.4) (Figure 7, [65])). The negative Pt/Pt* values (Pt/Pt* < 1) and low Pd/Ir
ratios are a characteristic feature of IPGE-elevated chromitites from Skyros and of low-PGE chromitites,in contrast to PPGE-elevated ones from the Pindos complex (Figure 7; Table 1).

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Minerals 2020, 10, 10

Figure 6. Plots of δ53 Cr values versus Cr/(Cr + Al) ratios and Ga for certain chromitites from the Balkan
Peninsula. There is a good negative trend between δ53 Cr values and Cr/(Cr + Al) ratios for chromitites
from Othrys and Skyros (a), and a positive trend between δ53 Cr values and Ga contents for chromitites
from Skyros and Vourinos–Othrys (b). Data from Table 1.

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Minerals 2020, 10, 10

Figure 7. Plots of Pd/Ir versus Pt/Pt* normalized values [(Pt/Pt* = (Pt/8.3)/(Rh/1.6) × (Pd/4.4)] after [65]of chromitities studied herein, grouped by locality. Green circle = Asthenosphere mantle, calculatedby [66]. Data from Table 1.

5. Discussion
Geochemical compositions and mineral chemistry characteristics of peridotites and chromitites
from Greece have shown that the partial melting degree of primitive mantle and the hydrous nature ofparent magma are major factors that control the compositional variations in them [10,18,19,42,45,56,57].
In addition, the known widespread geochemical heterogeneities in the bulk-rock compositions ofchromitites from the Balkan Peninsula (Table 1) would be inconsistent with their evolution throughsimple partial melting processes, and suggest that their original geochemical compositions may havebeen modiﬁed by subsequent, post-magmatic processes taking place at shallow mantle depths [8,66–69].
5.1. Potential Pathways Contributing to Chromium Isotope Fractionation
Chromium occurs in diﬀerent oxidation states in nature, with relatively immobile Cr(III) and very
soluble oxidized Cr(VI) being the most abundant species [70]. Redox processes are accompanied bysigniﬁcant isotope fractionation, also in systems where Cr(VI) is transformed to Cr(III) [71]. During
reduction, the lighter isotopes are preferentially reduced, resulting in an enrichment of 53 Cr relative
to 52 Cr values in the remaining Cr(VI) pool. This enrichment is measured as the change in the ratioof 53 Cr/52 Cr, and is expressed as δ53 Cr values in units permil (%) relative to a standard [32]. In thefollowing we consider potential pathways for changes in redox conditions that might aﬀect the Cr isotope
signatures and their changes in magmatic to post-magmatic processes: (1) measured δ53 Cr signaturesin mantle chromitites reﬂect primary magmatic compositional variations, (2) oxidative ﬂuids may aﬀect
solubility of Cr-bearing minerals and release of Cr(VI) during the formation of ferrian-chromite, (3)reductive ﬂuids may cause reduction of Cr(VI) back to Cr(III), deposition of graphite and Fe–Cr alloys
and (4) weathering of Cr-bearing minerals [19–21,25,60–64], and related dissolution/re-precipitation
processes potentially having a control on the bulk δ53 Cr signatures of chromitites. Based on theobservation that bulk silicate earth (BSE) carries a slightly lighter δ53 Cr signature compared to thatof mantle-derived chromitites, it has been suggested that small but discernible high-temperature Cr
isotope fractionations might occur within the Earth’s mantle, potentially imparted during fractional
crystallization and/or partial melting of mantle sources [1].
5.1.1. Magmatic Processes
It has been established that high-Cr and high-Al podiform chromitites are derived from melts
formed by high and low degrees of partial melting in the upper mantle, respectively [3,13,14,72].
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Minerals 2020, 10, 10

Considering the data distribution in the scatter-diagram of Pd/Ir versus the Pt/Pt* ratios for chromitites
from the Balkan Peninsula, it appears that chromitites from the large deposits (Vourinos and Othrys)
exhibit Pd/Ir < 1 and Pt/Pt* < 1, suggesting partial melting of already depleted mantle [65], and anincreasing partial melting trend from high-Al toward high-Cr deposits (Table 1; Figure 7). In addition,
the relatively high Pd/Ir ratio, that is considered to indicate the fractionation degree of the PGE [66],provides evidence for fractionation to some extent of parent magmas for the PPGE-elevated chromitites
from Pindos compared to the IPGE-elevated chromitites from Skyros and the Milia area of Pindos(Table 1; Figure 7).
The closer inspection of the Cr isotope data from chromitites studied herein, ranging from negative(−0.184%) to positive (+0.159%) δ53 Cr values (Table 1; Figures 5 and 6), may provide an opportunity
for a better understanding of the role of magmatic or/and post-magmatic processes on the Cr isotope
signatures of chromitites [2,21–23]. A comparison of the Cr isotope data for chromitites from our study(Table 1) with published values from other chromitites reveals a signiﬁcant overlap in the δ53 Cr data.
Our lowest δ53 Cr signature (−0.184%) is slightly higher than the most negatively fractionated δ53 Cr
value of −0.246% reported by [19], and the most negatively fractionated value of −0.21 % in the data
set published [1,2], while the isotopically most positively fractionated value in our data set (+0.159%)is somewhat higher than the hitherto highest reported δ53 Cr value (+0.068%) for global chromitites
from ophiolite complexes and layered intrusions [1,2].
The diagrams of δ53 Cr values versus Cr/(Cr + Al) ratios depict better deﬁned negative trends forchromitites from large deposits such as those from Othrys–Vourinos and Skyros (Figure 6a), than for
chromitites from smaller occurrences. Given that the former chromitites are characterized by small
degrees of fractionation of the respective parental melts, as exempliﬁed by the low Pd/Ir ratios (Table 1),
the observed correlation trends may point to the lowering of δ53 Cr values in the melts alongside withthe partial melting degree of the respective magma sources. In addition, the most evolved samples
from the lower cumulate sequence of the Vourinos ophiolite complex (sample 10) has higher Cr isotope
compositions, along with higher Fe contents, compared to the mantle chromitites (represented by
Voidolakkos, Kondro and Kissavos (Table 1), that may suggest a coupling of increasing trends of δ53 Cr
values with the degree of fractionation of the parent magma. Also, with respect to trace elements,
including Ni, Co, Zn, V, Mn, Ti and Ga, deﬁned correlations are only depicted between Ga and V
concentrations, and Cr/(Cr + Al) ratios. Negative trends are revealed by the Othrys–Vourinos andSkyros chromitites, both of which including high-Cr and high-Al ores (Figure 4). Considering thelow Pd/Ir ratios shown by those chromitites (Table 1), these negative trends could be attributed tocompositional variations in the parent magmas, rather than to changes inherent during fractionation ofthese melts.
5.1.2. Post-Magmatic Processes
In general, the lack of any clear correlation between major and trace elements, and δ53 Cr values
may be due to post-magmatic processes (dehydration, serpentinization, metamorphism), which may
modify the primary magmatic compositional trends recorded by chromite and silicates (depending on
the chromite/silicate ratio), during an extended period of deformation (ductile asthenospheric mantle
ﬂow to shallow crustal brittle deformation) [68,72–80]. With respect to the Rhodope massif, on thebasis of major-, minor- and trace-element compositions of chromite and Re–Os isotopic compositionsof magmatic PGM on the Dobromirtsi chromitites, a long-term history involving reworking of theancient Archean mantle, during the Phanerozoic, has been suggested [78–82]. Geochemical and mineralchemistry data demonstrated a complex interplay of substitutions, related to the ability of ﬂuids toinﬁltrate the chromite and the extent of the interaction between pre-existing cores and rims, and thatduring metamorphism minor and trace elements in chromite can be strongly modiﬁed [68]. Also,a sequence of petrological, deformation and redox events has been described for the upper-mantle
rocks of the Othrys ophiolite complex. These events are seen to have aﬀected the chromitite bodies
during their obduction, emplacement and subsequent alteration [36]. The compositional variations in

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Minerals 2020, 10, 10

the Skyros chromitites in particular [83], may have been caused by redox changes during multistagetransformations in the Othrys chromites. Cr oxidation and reduction processes accompanying thesedeformational events may have caused the variations of Cr isotope signatures measured in thosechromitites. (Table 1; Figures 4 and 6).
The variations of δ53 Cr signatures in the high-temperature magmatic environment in which
the chromitites from the Balkan peninsula were formed, seems to be in a good agreement withthe variations measured in some metamorphic Cr-rich minerals such as uvarovites, Cr-tremolite,
Cr-diopside, Cr-pyrope. Some of the positively fractionated δ53 Cr signatures of these phases havebeen interpreted as a consequence of heavy Cr isotope enrichments caused by metamorphism [2,21].
Moreover, the investigation of Cr isotopic composition of diﬀerent types of mantle xenoliths fromdiverse geological settings (fertile to refractory oﬀ-craton spinel and garnet peridotites, pyroxenite veins,metasomatised spinel lherzolites and associated basalts) revealed δ53 Cr values of peridotites which
also span a range from negatively to positively fractioned values of −0.51 to +0.75%, and a slightly
negative correlation between δ53 Cr and Al2 O3 and CaO contents for most mantle peridotites [23].
5.2. The Role of Sulphides on the δ53 Cr Signatures
The most positively fractionated δ53 Cr values are depicted by all massive chromitite samples
from the large Othrys complex, and by a few samples from small chromite occurrences (Table 1).
Chromitites in certain ophiolite complexes, such as Othrys, Lemesos, Cyprus, Shetland, Bulqiza,
Oregon, Moa-Baracoa and elsewhere are associated with Fe–Ni–Cu-sulphide mineralization withdominant minerals pyrrhotite, chalcopyrite and pentlandite, hosted in mantle serpentinized peridotites
and ultramaﬁc rocks of the cumulate sequence of the ophiolite complexes [9,17,52,53,62,63]. Although
an originally primary magmatic origin for these sulphides is not precluded, characteristics of the highly
transformed ore at the Eretria chromite deposit (Greece) may indicate that the original magmaticfeatures have been overprinted and that metals were released from the host rocks by a low-level
hydrothermal circulation process [52]. Therefore, in addition to a potential release of Cr from thechromite ore aﬀected by post-magmatic processes, such as brittle deformation and ﬂuid circulation,
the less alkaline environment which would be controlled by the dissolution of sulphides could be
potentially important for the control of redox-mediated Cr mobilizations and accompanying changesin δ53 Cr signatures.
5.3. The Potential Role of Abiotic Methane (CH4 )A salient feature of the presented data that potentially could explain the rather large compositional
and isotope variations, could be the presence of graphite (Figure 3), and the observation of methane [54]occurring in ﬂuid inclusions of fracture-ﬁlling minerals, along with the presence of chlorite andserpentine. These features are especially present in the Othrys chromitites, which also show the mostpositively fractionated δ53 Cr values (Table 1; Figure 5). Recently, the investigation of microdiamonds
and graphite in chromite from ophiolite-type chromitites hosted in the Tehuitzingo serpentinite(southern Mexico), have been attributed a secondary origin and process related to the retrograde
evolution of the respective chromitite. These processes apparently took place at relatively low
temperatures (520–670 ◦ C), at low pressures and in shallow depths [69]. The presence of minerals
indicative of super-reducing phases, such as graphite-like, Fe–Ni–Cr alloys, awaruite (Ni3 Fe), andheazlewoodite (Ni3 S2 ) [44,45,82,83], in the chromitite ores from the islands of Othrys and Skyros
may be suggestive of the introduction of continental crust derived, reduced C–O–H ﬂuids, during
post-magmatic alteration processes. These were probably facilitated during brittle deformation stages
in the shallow crust and in the subducted oceanic slab [69,84,85]. The association of graphite withsperrylite and sulphides in magnetite ore from the Skyros ophiolite complex has also been attributed
to its deposition from a CO2 − and CH4 − dominated reducing ﬂuid at low oxygen fugacity at relatively
low temperatures (500–300 ◦ C) during serpentinization [58]. In addition, methane is widespread in
microfractures and porous serpentine- or chlorite-ﬁlled veins [54]. The widespread occurrence of167

Minerals 2020, 10, 10

methane and Ni–V–Co phosphides in the Othrys chromitites [54,84,85] and graphite in the Skyrosisland ophiolite (Figure 3) seems to be compatible with the positively fractionated δ53 Cr values(Table 1), features pointing to the circulation of reducing ﬂuids, probably during serpentinization,
which facilitated the back-reduction of mobilized Cr(VI) fractions to secondary immobile Cr(III)-bearing
mineral phases [32,86].
Therefore, the relatively elevated, slightly heavy Cr isotope signatures recorded in chromitites
may be the result of a combination of (a) fractionation during magmatic processes, i.e., degree of partial
melting and diﬀerentiation of the parent magma, and (b) multi-stage post-magmatic overprinting,which seem to depend on the size of the respective chromitite occurrence (in terms of tonnage ofchromite) and on a probable relationship between the size and degree of shielding towards alterationof the chromitite bodies.
5.4. Weathering of Cr-Bearing Minerals and Environmental SigniﬁcanceContamination of groundwater and soil by heavy metals is becoming a serious threat to ourenvironments worldwide. The presence of harmful Cr(VI) in soils and in turn in groundwater may berelated to natural processes or/and human activities, such as transfer of weathered material from rocks
and primary raw materials, wastes and/or application of large amounts of fertilizers and pesticides
for a long time in cultivated areas. The European Soil Data Centre has established a link between theeﬀects of metal bio-availability and metal bio-accumulation, and human health and negative impacts
on ecosystems [25–30,87–96]. Dose-dependent diﬀerences in toxicities of elements, the particle size,
and the oxidation state require serious consideration in health-risk assessments [90]. A very common
source of salts in irrigated soils is the irrigation water itself, aﬀected by sea water in low-lying areas
along the coast. The salinization of shallow aquifers may have a major eﬀect on plant/crop growth
and toxicity. The use of multidisciplinary methods in the study of ecosystem processes in response to
groundwater and soil system has great potentials for sustainable developments [87–96].
The application of Cr isotope measurements to evaluate Cr(VI) contamination in groundwater androck leachates from the central Evia and Assopos basin in Greece has revealed positively fractionatedδ53 Cr values ranging from +0.56 to +0.96% in water leachates of ultramaﬁc rocks, which are comparable
to those from Central Europe [26–28]. These results imply oxidative mobilization of Cr(VI) fromthe ultramaﬁc host rocks, and successive back-reduction of so mobilized Cr(VI) fractions [29]. The
widespread distribution of ophiolite complexes in Europe and in orogenic zones elsewhere around theglobe, emphasizes the need for further more detailed studies addressing the variation of δ53 Cr valuesin Cr-bearing rocks and ores.
In summary, the original geochemical compositions of chromitites may have been modified bysubsequent, post-magmatic processes taking place at shallow mantle depths [35–38,56,57,68,69,72–89,97].
The wide range of δ53 Cr values, from positive, slightly negative to the most negatively fractionatedsignatures, even present in individual relatively large chromitite deposits, together with negative
trends observed between δ53 Cr values and Cr/(Cr + Al) ratios in chromitites, may reﬂect a control of
Cr isotope compositions in chromitites by degrees of partial melting and by magma fractionation. Thisis best exempliﬁed by the high-Al chromitites from the cumulate sequence of the Vourinos complex. In
addition, post-magmatic redox cycling of Cr-isotopes may occur in response to brittle deformation ofchromitites and the subsequent interaction of the Cr-bearing rocks with oxidative ﬂuids, also potentiallyleading to the formation of Fe-chromite through substitution of Cr, Al and Mg by Fe(III) and Fe(II). The
formation of Cr-bearing minerals, such as serpentine, chlorite, and some garnets, and the concomitantrelease of Cr(VI)-bearing ﬂuids, under alkaline conditions, may take place during serpentinization ofrespective ore bodies. Mobilized Cr(VI) bearing ﬂuids then are prone to back-reduction promotedby the reduced environment in serpentinized ultramaﬁc rocks, causing the formation of graphite,
secondary Fe–Cr-bearing silicates, and other secondary alteration phases that potentially can carry apositively fractionated Cr isotope signal.

168

Minerals 2020, 10, 10

6. Conclusions
The δ53 Cr values presented herein for chromitites from the Balkan Peninsula oﬀer an opportunity
to study the eﬀects of Cr isotope fractionations that potentially could be related to primary magmatic
processes and to compare these with compositional trends eventually depicted by the respective
chromite deposits and chromite occurrences.
A modiﬁcation of the magmatic control on the geochemical characteristics of chromitites is revealedby positively fractionated δ53 Cr values in ophiolite zones aﬀected by brittle deformation/metamorphism
and characterized by the presence of secondary epigenetic Cr-bearing minerals, formed by the circulationof reducing ﬂuids carrying abiotic methane.
Chromitites from the Balkan Peninsula depict a wide range in δ53 Cr values. Signatures range
from positively-, to slightly negatively fractionated δ53 Cr values, even in individual relatively large
deposits. Positively fractionated δ53 Cr values of all chromitite samples from Othrys and of high-Al
chromitites from Skyros, and a negatively correlated trend between δ53 Cr and Cr/(Cr + Al), may reﬂect
a control of δ53 Cr by degree of partial melting and by magma fractionation. This is best exempliﬁed byhigh-Al chromitites from the cumulate sequence of the Vourinos complex.
The application of Cr isotopes to the evaluation of Cr(VI) contamination in groundwater androck leachates relies on oxidative mobilization of Cr(VI) and successive back-reduction to C(III).
The widespread distribution of signiﬁcant natural Cr-sources in the Balkan peninsula, enhanced byindustrial activities involving Cr in their processes, present a latently dangerous framework which isprone to have negative eﬀects on the environment, particularly with respect to hexavalent Cr toxicity.
Detailed studies addressing the variation of δ53 Cr values in Cr-bearing rocks, ores, soils, surface run-oﬀ
and groundwater are required to understand the complicated mutual interplay between originalnatural Cr sources and the weathering-induced release of Cr into the environment.
Author Contributions: Conceptualization, R.F. and M.E.-E.; Methodology, M.E.-E., R.F. and I.M.; R.F. provided thechromium isotope analyses and their validation. M.E.-E. and I.M. provided the ﬁeld information and performed
the other analyses. All authors (M.E.-E., R.F. and I.M.) contributed to the elaboration, interpretation of the data
and to the writing of the manuscript. All authors have read and agreed to the published version of the manuscript.
Funding: This research was funded by the National and Kapodistrian University of Athens (NKUA) (Grant No.
KE_11730).
Acknowledgments: Many thanks are expressed to the two anonymous reviewers for their constructive criticism
and suggestions on an earlier draft of the manuscript. Vassilios Skounakis in thanked for his assistance with
SEM/EDS analysis of the samples. RF thanks Toby Leeper for always keeping the mass spectrometers at IGN inperfect running condition, and Toni Larsen for help in the ion chromatographic separation of the samples. This
paper is dedicated to the memory of Maria Zhelyaskova-Panayiotova, University of Soﬁa, who provided thechromitite samples from Bulgaria, and passed away in May 2018.
Conﬂicts of Interest: The authors declare no conﬂict of interest.

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174minerals
ArticleFactors Controlling the Gallium Preference inHigh-Al Chromitites
Ioannis-Porfyrios D. Eliopoulos * and George D. Eliopoulos
Department of Chemistry, University of Crete, Heraklion GR-70013, Crete, Greece; giorgoshliop@yahoo.gr* Correspondence: disaca007@hotmail.com; Tel.: +30-281-054-5136
Received: 6 August 2019; Accepted: 7 October 2019; Published: 10 October 2019

Abstract: Gallium (Ga) belongs to the group of critical metals and is of noticeable research interest.
Although Ga3+ is highly compatible in high-Al spinels a convincing explanation of the positive
Ga3+ –Al3+ correlation has not yet been proposed. In the present study, spinel-chemistry andgeochemical data of high-Al and high-Cr chromitites from Greece, Bulgaria and the Kempirsai Massif(Urals) reveals a strong negative correlation (R ranges from −0.95 to −0.98) between Cr/(Cr + Al)
ratio and Ga in large chromite deposits, suggesting that Ga hasn’t been aﬀected by re-equilibration
processes. In contrast, chromite occurrences of Pindos and Rhodope massifs show depletion in
Ga and Al and elevated Mn, Co, Zn and Fe contents, resulting in changes (sub-solidus reactions),
during the evolution of ophiolites. Application of literature experimental data shows an abrupt
increase of the inversion parameter (x) of spinels at high temperature, in which the highest values
correspond to low-Cr3+ samples. Therefore, key factors controlling the preference of Ga3+ in high-Al
chromitites may be the composition of the parent magma, temperature, redox conditions, the disorderdegree of spinels and the ability of Al3+ to occupy both octahedral and tetrahedral sites. In contrast,
the competing Cr3+ can occupy only octahedral sites (due to its electronic conﬁguration) and the Ga3+
shows a strong preference on tetrahedral sites.
Keywords: chromite; gallium; spinel; structure; composition; correlation; ophiolites; disorder

1. Introduction
Gallium (Ga) is a vital metal for the economy, due to its use in high-technology applications, suchas electronics industry, electric cars, solar panels. Although bauxite deposits are traditionally mined
for their Al content and are important sources of Ga as a byproduct commodity [1], the distributionof Ga in chromite ores may be of particular research interest, due to its relationship with the major
element composition of chromite.
Chromite belongs to the subgroup of spinels, which accommodate a wide variety of cationsin their structure with the general formula AB2 O4 . Many authors emphasized that despite theirsimple structure, many spinels exhibit complex disordering phenomena involving the two cationsites, which play an important role both in their thermochemical and their physical properties [2–6].
The movements of cations between tetrahedral and octahedral sites, as a result of cation substitution,
have been discussed under the aspect of structural parameters, such as tetrahedral and octahedralbond lengths, cation-cation and cation-anion distances, bond angles and hopping lengths, whichwere calculated by experimental lattice constants and oxygen parameters [2–4,7–12]. The ability ofGa3+ (r = 0.62 Å) to replace Al3+ (r = 0.54 Å) in aluminum minerals is related to their geochemistry(Group III of the periodic table), while Ga might be expected to behave in a similar way in chromite
and magnetite as they share similar ionic radii (Cr3+ , r = 0.62 Å; Fe3+ = 0.64 Å ) for octahedral andtetrahedral coordination (Al3+ , r = 0.39; Ga3+ = 0.47; Fe3+ = 0.49) [3]. Gallium levels reported in
Cr-spinel grains from ophiolites, varying from 10 to 50 ppm [13,14] are consistent with experimental

Minerals 2019, 9, 623; doi:10.3390/min9100623

175

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Minerals 2019, 9, 623

mineral-melt data on the partition coeﬃcient (DGa = 0.9–11.2) [15]. Although mineral–melt partitioncoeﬃcients are not constants, depending on a number of factors (pressure, temperature, oxygen
fugacity or mineral and melt composition) on the basis of experimental data it has been suggested that
Ga is volatile and there is no signiﬁcant eﬀect of temperature, magma composition and at very low
oxygen fugacity conditions [15].
A tectonic discrimination of peridotites, using the oxygen fugacity (f O2 )–Cr#[Cr/(Cr + Al)] diagram,and the Ga–Ti–Fe3+ # [Fe3+ /(Fe3+ + Cr + Al)] systematics in chrome-spinels, has been proposed [16].
A negative correlation between Ga and Cr in chromitites has been established, that may be related tothe composition of parental magmas [17–24] or to the outer electronic structure of Ga that is similar tothat of Al [25]. Also, the investigation of spinels in lithospheric mantle xenoliths from distinct tectonic
settings has demonstrated that trace elements contribute in discriminating between spinels hosted
in peridotites and those crystallized from magmas [11]. However, a convincing explanation of thepositive correlation between Ga3+ and Al3+ has not yet been oﬀered.
In the present study we characterize the spinel chemistry, bulk ore composition, including Ga,
from chromitite samples of selected ophiolite complexes in Greece (Pindos, Central Vourinos andSkyros), the Rhodope–Serbo–Macedonian zone (SMZ) massifs (all of Mesozoic age) and the Kempirsai
Massif (Kazakhstan) in the Urals (Palaeozoic age), all hosting both high-Cr and high-Al chromitites.
The investigated samples are representative of large chromite deposits and small occurrences, inorder to deﬁne potential relationships between major, minor or trace elements and Ga and theeﬀect of re-equilibration processes, during a long evolutionary time of the ophiolites. We apply
available platinum-group element (PGE) data to deﬁne potential correlations between the Ga content
and fractional crystallization, and experimental literature data for the structure of spinels, aiming
to investigate the role of intra-crystalline cation exchange, and contribute to still uncertain factors
controlling the positive Al-Ga correlation of chromitites.
2. Materials and Methods
2.1. Mineral Analysis
Polished sections of all chromitite samples were examined using a reﬂected light microscope
and a scanning electron microscope (SEM), equipped with energy-dispersive spectroscopy (EDS).
The SEM-EDS back-scattered electron images (Figure 1) and analyses of chromite ores (Table 1) werecarried out at the Faculty of Geology and Geoenvironment, National and Kapodistrian Universityof Athens (NKUA), using a JEOL JSM 5600 (Tokyo, Japan), scanning electron microscope, equippedwith ISIS 300 OXFORD (Oxford shire, UK), automated energy dispersive analysis system. Analyticalconditions were 20 kV accelerating voltage, 0.5 nA beam current, <2 μm beam diameter and 50 s counttimes. The following X-ray lines were used: FeKα, NiKα, CoKα, CuKα, CrKα, AlKα, TiKα, CaKα,SiKα, MnKα and MgKα. Cr, Fe, Mn, Ni, Co, Ti and Si, MgO for Mg and Al2 O3 for Al. Contents of
Fe2 O3 and FeO were calculated on the basis of the spinel stoichiometry.

176

99.5

100.3

0.77
0.65
0.031

0.019
3.426
12.037
5.170
2.744
0.498
0.065
0.157

Total

Cr/(Cr+Al)
Mg/(Mg+Fe2+ )Fe3+ /(Cr+Al+Fe3+ )

Ti
Al
Cr
Mg
Fe2+
Fe3+
Mn
Ni

177

Sky. 4

0.1
11.2
59.1
13.8
12.1
2.4
0.2

wt%

TiO2
Al2 O3Cr2 O3MgO
FeO
Fe2 O3
MnO

0.2
12.6
58.1
14.4
11.5
2.2
0.2

Sky. 5

Skyros

0.039
2.889
13.031
4.703
3.271
0.001
0.044
0.021

0.81
0.62
0.000

0.2
9.4
63.2
12.1
14.1
0.1
0.2
0.1

0.3
11.4
59.7
13.6
13.1
1.7
0.3
0.2

TiO2
Al2 O3Cr2 O3MgO
FeO
Fe2 O3
MnO
NiO

Vour. 2

Vour. 1wt%

Vourinos

n.d.
15.8
53.9
13.4
13.4
3.1
0.2

Souﬂi 1

0.040
2.997
12.523
5.250
2.744
0.400
0.045
0.000

0.81
0.66
0.025

99.9

0.2
9.8
61.3
13.6
12.7
2.1
0.2
n.d.

Vour. 3

0.2
19.6
47.8
14.1
13.5
5.0
0.2

99.9

0.2
26.2
41.5
14.5
13.5
3.4
0.3
0.3

Pi.1.

99.4

0.3
27.2
39.5
13.6
15.7
3.1
n.d.
n.d.

Pi.2

100.1

0.2
32.9
35.9
16.4
12.3
2.3
0.1
0.2

Pi. 3

101.2

0.2
34.5
34.8
16.7
12.5
2.6
n.d.
n.d.

Pi. 4

Pindos

0.2
30.8
35.5
15.3
13.9
5.1
n.d.

Gomati
0.4
27.9
37.3
16.6
10.4
6.0
0.2

Broucevci
0.2
5.5
58.8
10.2
16.2
7.2
0.6

Jacovitsa

0.5
10.4
49.8
8.2
20.7
9.5
0.5

Pletena

Bulgaria

0.019
4.650
10.645
4.712
3.222
0.659
0.043
0.041

0.70
0.59
0.0413

100.7

0.1
16.5
52.8
12.4
15.6
2.9
0.2
0.2

Pi. 5

0.039
4.728
10.662
3.896
4.056
0.533
0.065
0.020

0.69
0.49
0.028

99.7

0.2
15.5
52.1
10.1
18.7
2.7
0.3
0.1

Pi.6

0.4
4.5
33.2
11.9
13.6
35.1
n.d.

0.3
1.2
24.9
7.2
20.2
45.6
0.2

0.3
24.4
46.6
14.3
14.1
0.3
n.d.

0.3
23.3
47.1
14.2
14.1
0.8
n.d.

0.037
5.896
9.602
5.156
2.840
0.427
0.021
0.020

0.62
0.64
0.027

99.01

0.2
20.11
48.8
13.9
13.6
2.2
0.1
0.1

Sky. 1

0.036
6.922
8.586
5.610
2.367
0.420
0.020
0.039

0.55
0.69
0.026

100.1

0.2
24.5
45.3
15.7
11.8
2.3
0.1
0.2

Sky. 2

Skyros

0.1
9.3
61.3
15.5
9.9
4.5
n.d.

0.2
9.8
60.2
15.2
10.3
4.3
n.d.

Main Ore Field

Southern Part

Urals Kempirsai

0.020
2.555
12.835
4.611
3.300
0.570
0.067
0.042

0.81
0.58
0.036

99.6

0.1
8.2
61.4
11.7
14.9
2.8
0.3
0.2

Pi.7

Goliamo
Northern Part
Kamenyane Kamenyane
Batamshinsk
1
2

0.78
0.52
0.49
0.42
0.4
0.60
0.65
0.61
0.71
0.72
0.0055
0.038
0.036
0.025
0.027
Numbers of Cations on the Basis of 32 Oxygens
0.039
0.036
0.054
0.035
0.014
3.451
7.415
7.758
8.966
9.241
12.383
7.891
7.558
6.563
6.254
4.790
5.210
4.906
5.652
5.658
3.164
2.706
3.149
2.326
2.377
0.087
0.620
0.575
0.401
0.436
0.044
0.061
0.000
0.020
0.000
0.042
0.058
0.000
0.037
0.000

99.8

0.2
11.3
60.4
12.4
14.6
.5
0.2
0.2

Vour. 5

Rhodope Massif

Souﬂi 2

Greece

0.018
6.736
8.471
5.774
2.203
0.757
0.040
0.000

0.56
0.72
0.036

100.5

0.1
23.9
44.8
16.2
11.1
4.2
0.2
n.d.

Vour. 4

Table 1. Electron scanning electron microscope (SEM)/energy-dispersive spectroscopy (EDS) analyses of chromite from chromitites of Greece, Bulgaria and Kempirsai(Urals).

0.038
6.532
9.340
3.813
4.142
0.053
0.042
0.040

0.58
0.51
0.0033

100

0.2
22.1
47.1
10.2
19.7
0.3
0.2
0.2

Sky. 3Minerals 2019, 9, 623

0.75
0.68
0.03

0.019
3.414
12.066
5.321
2.613
0.480
0.044
0.041

Cr/(Cr+Al)
Mg/(Mg+Fe2+ )Fe3+ /(Cr+Al+Fe3+ )

Ti
Al
Cr
Mg
Fe2+
Fe3+
Mn
Ni

0.038
3.813
11.695
5.512
2.442
0.414
0.044
0.041

0.79
0.69
0.026

99.7

0.2

0.2

99.1

NiO

Sky. 5

Sky. 4

Total

wt%

Skyros

0.019
4.636
10.745
5.036
2.940
0.580
0.043
0.000

0.70
0.62
0.036

100.2

n.d.

Souﬂi 1
0.2

0.036
5.591
9.396
5.231
2.805
0.940
0.000
0.000

0.63
0.65
0.059

99.8

99.1

0.2

Broucevci

99.1

0.2

Jacovitsa

100

0.2

Pletena

Bulgaria

0.44
0.47
0.88
0.74
0.66
0.72
0.52
0.37
0.057
0.068
0.093
0.123
Numbers of cations on the basis of 32 oxygens
0.035
0.035
0.041
0.102
8.475
7.828
1.767
3.277
6.553
7.021
12.67010.582
5.324
5.890
4.143
3.291
2.711
2.067
3.715
4.653
0.901
1.079
1.481
1.936
0.000
0.040
0.138
0.115
0.000
0.038
0.044
0.044

100.3

n.d.

Gomati

Rhodope Massif

Souﬂi 2

GreeceTable 1. Cont.

0.082
1.452
7.168
4.858
3.114
7.215
0.000
0.114

0.83
0.52
0.455

99.2

0.3

0.062
0.386
5.648
3.088
4.839
9,844
0.046
0.088

0.94
0.38
0.62

99.9

0.4

0.054
6.987
8.880
5.137
2.879
0.024
0.000
0.039

0.56
0.64
0.0015

100.4

0.2

0.054
6.,675
9.070
5.151
2.864
0.146
0.000
0.038

0.58
0.65
0.009

99.9

0.2

Goliamo
Northern Part
Kamenyane Kamenyane
Batamshinsk
1
2

0.019
2.787
12.323
5.874
2.125
0.852
0.000
0.020

0.82
0.74
0.053

100.8

0.1

0.038
2.947
12.143
5.780
2.197
0.833
0.000
0.062

0.80
0.73
0.052

100.3

0.3

Main Ore Field

Southern Part

Urals Kempirsai

Minerals 2019, 9, 623

178

Minerals 2019, 9, 623

Figure 1. Representative back-scattered electron images of chromite ores from Greece, showing texture
relationships between chromite and silicates, the presence of homogeneous chromite (a) and alsoabundant silicate inclusions (chlorite and serpentine) in the host chromite (b), porous texture andalteration to Fe-chromite (c,d). Abbreviations: chr = chromite; Fe-chr = iron-chromite; srp = serpentine;
chl = chlorite.

2.2. Whole Rock Analysis
Major and trace elements in massive chromitite samples (more than 95 vol %) were determined by
ICP-MS analysis, at the ACME Analytical Laboratories Ltd., Vancouver, BC, Canada (currently Bureau
Veritas Commodities Canada Ltd.). The samples were dissolved using an acid mix (HNO3 –HClO4 –HF)
digestion and then the residues were dissolved in concentrated HCl. The rare earth elements La, Ce,
Pr, Nd, Sm, Eu, Gd, Tb, Dy, Ho, Er, Tm, Yb and Lu as well as Li, K, Ge, Sr, Y, Zr, Mo, Sb, Cs, W, Pb,
Th, and U were lower than the detection limits of the analytical methods. The detection limits of themethod for the presented elements are 1 ppm for Ga and V, 0.2 ppm for Co and Zn, 0.1 for Ni and0.01 wt % for Fe. On the basis of the quality control report provided by the Analytical Labs, the resultsof analyses of the reference material in comparison to expected values, and the results from multistage
analysis of certain samples, showed accuracy and precision in good agreement with accepted values
for international standards. The analytical error, for Ga, for example, was <5%. Although the PGE
data reported in Table 2 have been published previously (Table 2) a portion from the same samples
was used for the presented trace element analyses.

179

180

0.69

0.54

0.72
0.52
0.37
0.52
0.38

0.47
0.88
0.74
0.83
0.94
0.80
0.82
0.56
0.64

Kemprsai (Urals)
Northern
Batamshinsk
Southern
XL Let Kazakhstan

1600
1600
1500
1700

1300
2250
890
1550
2260

1700
1150
1030

1400

1300
1600
1500
1250
1200

1300
1500
1500
1630
750
720
1450

2000
1580
1900
1800
1800

Ni

210
230
240
230

230
310
290
80
970

230
220
130

210

250
200
240
220
200

140
260
260
240
270
240
290

240
210
200
170
180

160
200
680
730

790
240
330
1000
370

380
460
730

960

1200
1000
870
640
620

600
710
710
580
760
760
560

500
560
400
780
620

V

160
190
480
340

420
760
1030
450
4030

580
280
280

370

540
400
450
420
400

400
410
520
460
520
620
490

260
550
300
360
280

Zn

Trace Element (ppm)Co

Symbol * = Data on PGE from literature [26–32].

0.73
0.74
0.58
0.65

0.62
0.65
0.67

0.70
0.63
0.49

Rhodope Massif
GreeceSouﬂi1
Souﬂi2
Gomati
Bulgaria
Broucevci
Jacovitsa
Pletena
Goliamo Kamenyane 1
Goliamo Kamenyane 2

Othrys (n = 4)

0.64
0.69
0.51
0.69
0.69

0.61
0.55
0.58
0.75
0.79

Skyros 1
Skyros 2
Skyros 3
Skyros 4
Skyros 5

0.63
0.6
0.71
0.72
0.57
0.49
0.58

0.52
0.48
0.42
0.4
0.67
0.69
0.81

Pindos 1
Pindos 2
Pindos 3
Pindos 4
Pindos5
Pindos 6
Pindos 7

0.65
0.62
0.66
0.72
0.61

0.77
0.81
0.81
0.56
0.78

Mg/(Mg+Fe2+ )SEM/EDS
Cr/(Cr+Al)

Vourinos 1
Vourinos 2
Vourinos 3
Vourinos 4
Vourinos 5

Location

14
16
48
49

45
9
13
12
6

16
12
24

33

34
36
40
14
11

23
32
34
32
16
23
17

14
15
13
27
14

Ga

Table 2. Trace element contents of high-Cr and high-Al chromitites.

9.1
9.5
12.6
10.1

11
13.9
17.9
11.6
64.3

13.2
10.6
9.6

9.8

10.9
9.45
9.6
10.7
10.1

7.84
9.77
7.8
7.3
9.7
10.1
10.2

8.1
7.9
8.72
9.2
7.76

Fe

wt%

ppb

60
197
563
87
40

150
82
104

91

2300
464
251
145
145

51
143
117
6123
3875
2098
181

140
135
92
30
109

ΣPGE *

0.92
0.46
0.07
0.14
1.93

0.2
1.0
0.25

0.36

0.08
0.7
0.67
0.33
0.1

1.0
1.0
6.33
34.2
12.2
7.2
0.22

0.06
0.35
0.55
0.54
0.63

Pd/Ir *

Minerals 2019, 9, 623

Minerals 2019, 9, 623

3. A Brief Outline of Characteristics for the Studied Chromitites
All chromitite samples selected for the present study come from deposits and occurrences, whichhave been the subject of detailed geological, mineralogical and geochemical investigation [26–36] andreferences therein. The main ophiolite complexes of Greece (Vourinos, Othrys and Pindos) belong to
the Upper Jurassic to Lower Cretaceous Tethyan ophiolite belt, and are characterized by heterogeneous
deformation and rotation, during their original displacement and subsequent tectonic incorporation
into continental margins [33]. The studied samples of chromitites are massive (Figure 1) and exhibit
variations in the chromitite tonnage, the composition of chromite (Tables 1 and 2), the degree oftransformation of ores and the associated ophiolites [27–37].
Chromite ores in the Vourinos complex occur in the mantle and cumulate sequences, with atonnage estimated to approximately 10 Mt of high-Cr type, but at the central part of the complex there
are high-Cr and high-Al ores in a spatial association, with low PGE contents [26]. The Othrys complex
has a relatively high tonnage (approximately 3 Mt) of high-Al massive chromite ores and low PGE
content [27].
The chromitite occurrences in the Pindos ophiolite complex are small (a few tens of m (x) a few
tens of cm) and are hosted within completely serpentinized and weathered, intensively deformed
dunite-harzburgite blocks, due to a strong plastic and brittle deformation that was superimposedon primary magmatic textures [17,18,34]. Chromitites throughout the Pindos complex are high-Cr
and high-Al, often in a spatial association. The most salient feature of the Pindos chromitites is theenrichment in Pt and Pd at the area of Korydallos, at a level of 7 ppm PGEtotal [28,29] and up to
29 ppm [35]. In the Achladones area on the Skyros island small massive chromitite bodies are ofhigh-Al type and have elevated PGE contents, up to 3 ppm ΣPGE, although both high-Cr and high-Al
types having low PGE content are found on the entire island [30].
Ophiolites associated with the Serbomacedonian massif (Gomati) and Rhodope massif including
the ophiolites of Souﬂi (Greece), Dobromirci, Jacovitsa, Broucevci and Goliamo-Kamenjane (Bulgaria)
host small (a few thousand tons) high-Cr and high-Al ores in a spatial are association, whichoccasionally contain elevated PGE concentrations. They are completely serpentinized, locally sheared
and metamorphosed to antigorite-tremolite and/or talc schists. Detailed description of the characteristic
mineralogy and texture of those chromitites have been published in previous studies [17,18,31,32,36].
The Kempirsai massif, covering an area of 2000 km2 , is divided by a shear zone into two parts:
the southeastern part that is called Main Ore Field (MOF), hosting large high-Cr chromite deposits,
and the northwestern area, the so-called Batamshinsk Ore Field (BOF), hosting much smaller high-Al
chromite deposits [37]. An excellent description of the petrography and mineral chemistry, including
mineral inclusions in the chromite of the giant chromite deposit of Kempirsai has been provided anddiscussed by Melcher et al [37]. These authors have interpreted their formation by a multistage process:High-Al chromitites may be derived from MORB-type tholeiitic melts, and high-Cr ones from boninitic
magmas, during a second stage by interaction of hydrous high-Mg melts and ﬂuids with depleted
mantle in a supra-subduction zone setting.
4. Results
4.1. Compositional Variations in Chromite
The chromite samples from the central part of the Vourinos complex, the Pindos, Skyros island,
Serbomacedonian, Rhodope and Kempirsai massifs show a wide variation in major elements fromhigh-Cr, with the Cr/(Cr + Al) atomic ratio ranging from 0.81 to 0.69, to high-Al with the Cr/Cr +Al) ratio ranging from 0.63 to 0.4 (Table 1), falling in the range of metallurgical and refractorytype,respectively [21]. In addition, in the Bulgarian Rhodope massif (Jacovitsa, Pletena and Goliamo
Kamenyane areas) altered chromite grains are dominant, having relatively high FeO and low Al2 O3
and MgO contents (Table 1). As a consequence, the Cr/(Cr + Al) atomic ratios of those chromitites are
signiﬁcantly higher than those of high-Cr chromitites from the Vourinos complex (Table 1).

181Minerals 2019, 9, 623

4.2. Distribution of Trace Elements in Chromitites
The geochemical data from whole rock analyses show a wide variation in major and trace element
contents (Table 2). Gallium contents are lower in high-Cr chromitites (11 to 23 ppm) compared tohigh-Al ones (27–49 ppm), that seems to be independent on the degree of fractionation of parent
magma, as exempliﬁed by the Pd/Ir ratio [38]. The highest Co, Mn, Zn, Fe and lowest Ga were mainly
recorded in strongly altered small chromite occurrences from the Rhodope massif in Bulgaria. They arein a good agreement with other chromitites [21–25] and are independent of the age of the associatedophiolites. Platinum-group elements (PGE) show total contents ranging from 30 to 6120 ppb and Pd/Ir
ratios from 0.06 to 34, which are independent of the major element composition of chromitites (Table 2).
The results show a strong negative correlation (R ranges from −0.98 to −0.95) between the Cr/(Cr+ Al) atomic ratio and Ga for the relatively large chromite deposits of Vourinos, Kempirsai massif(Urals) and the Skyros island. In addition, there is a less strong negative correlation for small chromite
occurrences from the Pindos and Rhodope massifs (R ≥ −0.76 and R ≥ −0.83, respectively). Apartfrom Ga, the best correlation is found between Cr/(Cr + Al) and V for the Vourinos (R = −0.84), Skyros
and Kempirsai (R ≥ −0.93), whereas no signiﬁcant relationship for chromitites from the Pindos andRhodope massifs (Figure 2b) or between Cr/(Cr + Al) and other minor and trace elements is observed.

Figure 2. Cont.

182

Minerals 2019, 9, 623

Figure 2. Plots of the Cr/(Cr + AI) atomic ratio versus Ga content (a); the Cr/(Cr + AI) ratio versus V
content (b) and the Fe3+ /(Al + Cr + Fe3+ ) ratio versus Ga content (c), including a detailed plot for the
Vourinos and Kempirsai samples (d). Data from the Table 2.

5. Discussion
5.1. Factors Controlling the Spinel Chemistry
5.1.1. Magmatic Versus Post-Magmatic Processes
The wide variation of the Cr/(Cr + Al) atomic ratio for the chromitite samples from Greece,
Bulgaria and Kempirsai massifs (Tables 1 and 2) fall in the range of metallurgical and refractory types.
Diﬀerences in the trace element content (Table 2) may suggest trace element partitioning, depending
on a number of factors, such as temperature, pressure, oxygen fugacity and the chemical compositionof parent magmas [17–47].
As the partition coeﬃcient of elements is deﬁned as the ratio of the element content in a mineraland the melt [15], the Pd/Ir ratio can be used as an indicator of the degree of fractionation of parent
magma for chromitites [38] and the presence of very low Pd/Ir values (low degree of fractionation) for
both high-Cr and high-Al chromitites, suggest their origin from diﬀerent magma sources. It has beenargued that high-Cr chromitites, which have higher Sc, Mn, Co and Ni, and lower Ti, V, Zn and Ga
contents may be derived from boninitic magmas, while high-Al ores may be derived from MORB-type
tholeiitic magmas [17–24]. Experimental data at high temperature have shown that Ga is compatible inspinel with D values ranging between 0.9 and 11.2, and slightly lower D values in the most reducing
experiments, while experimental data at temperatures >1300 ◦ C and low oxygen fugacity have shownthat there is no signiﬁcant eﬀect of temperature, composition and redox conditions [15]. However,the negative correlation between the Cr/(Cr + Al) atomic ratio and Ga content in natural chromitites
points to the potential eﬀect of the composition of the parent magma, while a positive trend between
the Fe3+ /(Al + Cr + Fe3+ ) atomic ratio and Ga content for large chromite deposits (Figure 2c,d) may
suggest the eﬀect of the redox conditions on the Ga distribution in chromitites. Speciﬁcally, high-Cr
chromitites formed earlier from a primary magma (under relatively reducing conditions), compared tohigh-Al ones formed later from an evolved magma (and more oxidized conditions) magma [47].
In addition, diﬀerences in the negative correlations between the Cr/(Cr + Al) atomic ratio and Ga
content and the slope of correlation lines for the diﬀerent occurrences (Figure 2) may suggest that inaddition to the composition of parent magmas, which is a major factor for large deposits (like Vourinos
and Kempirsai massif, Urals) other factors such as temperature, pressure or redox conditions may
be responsible for the observed deviation from linearity for small metamorphosed occurrences ofchromitites, such as those from the Pindos and Rhodope massifs (Figure 2b). The lack of signiﬁcantrelationships between major and trace elements, in small chromitite occurrences from the Pindos

183Minerals 2019, 9, 623

and Rhodope massifs (Figure 2) may be related with post-magmatic processes. The elevated Mn,
Co, Zn and Fe contents and depletion in Ga (Table 2) is consistent with the spinel chemistry in the
Rhodope massif of Bulgaria, showing a trend of depletion in Ga, in the metamorphic Fe-chromite rims
surrounding the cores of chromite grains, implying that most tetrahedral sites are still occupied by
Fe2+ [17]. In addition, it has been suggested that the Mg cations can be replaced by Mn, Zn or Co,
whereas Al and Fe3+ compete for the octahedral sites, hampering the entry of Ga [17,18,24].
Despite the recorded modiﬁcation in trace elements by re-distribution during post-magmatic
processes, as exempliﬁed by bulk analysis (Tables 1 and 2) and spinel chemistry [17,18,24] limited only
to relatively small chromitite occurrences, the well-established relationship (R = −0.95 to −0.98) between
Cr/(Cr + Al) ratio and Ga (Table 2; Figure 1a) that is comparable to literature data for chromitites hosted
in other ophiolite complexes [21–24,40–47] seems to be a salient feature.
5.1.2. Spinel Structure
The structure of spinel is a cubic close-packed array of 32 oxygen ions, with 64 tetrahedral
vacancies and 32 octahedral vacancies in one unit cell each, containing 8 formula units, with the generalformula: A2+ B3+ 2 O4 , where A = Fe2+ , Mn2+ , Mg2+ , Co2+ , Zn2+ , Ni2+ and B = Fe3+ , Cr3+ , Al3+ , Ga3+ ,V3+ [36]. Spinels are traditionally denoted as either “normal”, where the A cation occupies T sites,
the B cation occupies M sites, whereas in the “inverse” type cation B occupies the T site and the Msite is occupied by both cations A and B [4,8–12]. The degree of inversion x characterizing the cation
distribution can show values between x = 0 (normal spinel) and x = 1 (inverse spinel). The spinel
structure is able to accommodate many cations (at least 36) by enlarging and decreasing its tetrahedral
and octahedral bond distances, while the oxygen positional parameter (u) should be regarded as ameasure of distortion of the spinel structure from cubic close packing or as the angular distortion of
the octahedron [10,12]. The movements of cations between tetrahedral T and octahedral M sites, as aresult of Mg2+ substitution, can be discussed based on structural parameters, such as bond lengths,
cation-cation and cation-anion distances, bond angles and hopping lengths, which were calculatedusing experimental lattice constants and oxygen parameters [4,8–12].
5.1.3. Applications to Natural Spinels
Despite post-magmatic compositional changes in Fe-chromite within the chromitite ores, resulting
in the remobilization of cations during metamorphism (700 ◦ C to 450 ◦ C) of chromitites [17,18,21,24,32],
the structural incorporation of Ga into the chromite lattice is evidenced by the progressive and linear
increase of Al or decreasing Cr/(Cr + Al) atomic ratio (Figure 2a). Experimental data on cation
distribution versus temperature may provide valuable information related to the preferences of cationsin the spinel lattice. MgAl2 O4 is the most prominent example of a normal spinel because Mg2+ is muchlarger than Al3+ [2,6,48]. Gallium (Ga3+ ) is smaller than Mg2+ , but signiﬁcantly larger compared toAl3+ , leading to an ordering which is called mainly inverse, at least for the end-member composition
MgGa2 O4 [4].
The intra-crystalline exchange reaction in spinels has been modeled [49,50] and the order-disorderprocess has been described by the following exchange reaction: T Al + M Mg = T Mg + M Al (where
T = tetrahedral and M = octahedral site) in which the forward reaction implies an exchange of
Mg with Al at the M site (ordering process), and backwards (disordering process). The cation
distributions for both disordering and ordering experiments were obtained by measuring the oxygen
positional parameter (u), the inversion parameter (x) (Al in T) site, using samples with varying
composition. The Mueller kinetic model was satisfactory applied to the experimental data andallowed the calculation of the kinetic ordering constants K, linearly related to temperature by means ofArrhenius equations [48–51].
Martignago et al. [52] performed crystal structure reﬁnements on three natural spinels, on low-Cr
spinel containing small Fe3+ quantities and two other samples with high Cr (8.4 wt % Cr2 O3 ) and
low Cr (3.3 wt % Cr2 O3 ) contents [52]. These experiments have shown that both parameters (u) and184

Minerals 2019, 9, 623(x) remained constant for the three diﬀerent samples up to 600 ◦ C, independently of Cr3+ contents.
The distortion of spinels, started at higher temperature, near to 650 ◦ C (Figure 3). The degree ofdistortion at the highest temperatures is inversely correlated with Cr3+ contents. Since Cr3+ has atendency to be completely partitioned on the octahedral site, due to its electronic conﬁguration andsize [2,52–54], Al3+ cation is unable to substitute Cr3+ . The most salient feature derived from the above
experimental data [52] is the abrupt increase of the inversion parameter (x) having the highest values
for the sample with the lowest Cr content (L-Cr sample), and the lowest values for the sample with thehighest Cr content (H-Cr) [52].
Such structural changes should cause modiﬁcations of the structural, physical and thermal
properties of the spinels. Although the spinel structure is complicated and determination of several
parameters may be required, the above results [52], which show that the Cr content in spinels aﬀects theoccupancy of Al in the tetrahedral site, may suggest that the co-existence of Ga3+ in high-Al chromitites(Table 2) [2,4–12] is related to the degree of disorder that is inversely correlated with Cr contents.

Figure 3. Plot of the Inversion parameter (x) versus temperature for spinels after [52]. A potential trend
for chromitites with much higher Cr content (Table 1) is presented by the green arrow.

Therefore, potential controlling factors on the Ga preference in high-Al chromitites are (a) thecomposition of the parent magma (geotectonic setting), temperature and redox conditions, (b) the
electronic conﬁguration of Cr3+ resulting in occupation of M sites only, the ability of Al3+ to occupy
T and M sites and the strong preference of Ga3+ to T sites, and (c) elevated values of the inversion
parameter (x) that is inversely correlated with Cr3+ content and favors the co-existence of Ga3+ in
high-Al chromitites.

185

Minerals 2019, 9, 623

6. Conclusions
The presented geochemical and mineral chemistry data on chromitites associated with ophiolitecomplexes, in conjunction with experimental literature data allowed us to draw the following
conclusions:
(1)

(2)
(3)

(4)

(5)

The lower Ga contents in high-Cr chromitites (11 to 23 ppm) compared to high-Al ones (27–49 ppm)
suggest that the composition of the parent magma may be a major factor controlling the preferenceof Ga in high-Al chromitites.
The positive trend between the Fe3+ /(Al + Cr + Fe3+ ) atomic ratio and Ga content for large chromite
deposits may suggest the eﬀect of the redox conditions on the Ga distribution in chromitites.
Plot of the Cr/(Cr + Al) atomic ratios versus Ga content exhibits diﬀerences in terms of the slopeof correlation lines for the diﬀerent occurrences, suggesting that, in addition to the compositionof parent magmas, other factors such as temperature, pressure or redox conditions may aﬀect theobserved deviation from linearity for small metamorphosed chromitite bodies.
The depletion of Ga and Al, and elevated Mn, Co, Zn and Fe contents in certain small chromitite
occurrences, transformed during post-magmatic metamorphism, suggest potential change of the
Ga content in Cr-spinel during sub-solidus reactions.
Assuming that low-Cr spinel is characterized by the highest value of the inversion parameter (x)at higher than 650 ◦ C, then the high Al content in spinels may be a driving force for the degree of
inversion in the structure that facilitate the substitution of Al3+ for Ga3+ at magmatic conditions.

Author Contributions: I.-P.D.E. provided the SEM/EDS analyses. Both authors I.-P.D.E. and G.D.E. contributed tothe elaboration and interpretation of the data, and carried out the ﬁnal revision of the manuscript.
Funding: This research was not funded because it is mostly based on valuable literature data.
Acknowledgments: Many thanks are expressed to the National University of Athens for the donations thechromitite samples and the access to the analytical facilities (electron microprobe analysis). Many thanks are
expressed to Heinz-Gunter Stosch, Karlsruhe Institute of Technology (KIT), Institute for Applied Geosciences,
Germany, Maria Perraki, Technical University of Athens and the anonymous reviewers for the constructive
criticism and suggestions on an earlier draft of the manuscript. The linguistic improvement of this work by H.G.
Stosch is greatly appreciated.
Conﬂicts of Interest: The authors declare no conﬂict of interest.

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minerals
ArticleTrace Element Distribution in Magnetite Separates ofVarying Origin: Genetic and Exploration SigniﬁcanceDemetrios G. Eliopoulos 1 and Maria Economou-Eliopoulos 2, *1
2*Institute of Geology and Mineral Exploration (IGME), Sp. Loui 1, Olympic Village, GR-13677 Acharnai,
Greece; eliopoulos@igme.gr
Department of Geology and Geoenvironment, University of Athens, 15784 Athens, Greece
Correspondence: econom@geol.uoa.gr

Received: 10 November 2019; Accepted: 2 December 2019; Published: 6 December 2019

Abstract: Magnetite is a widespread mineral, as disseminated or massive ore. Representative
magnetite samples separated from various geotectonic settings and rock-types, such as calc-alkaline
and ophiolitic rocks, porphyry-Cu deposit, skarn-type, ultramaﬁc lavas, black coastal sands, andmetamorphosed Fe–Ni-laterites deposits, were investigated using SEM/EDS and ICP-MS analysis.
The aim of this study was to establish potential relationships between composition, physico/chemical
conditions, magnetite origin, and exploration for ore deposits. Trace elements, hosted either in themagnetite structure or as inclusions and co-existing mineral, revealed diﬀerences between magnetite
separates of magmatic and hydrothermal origin, and hydrothermal magnetite separates associatedwith calc-alkaline rocks and ophiolites. First data on magnetite separates from coastal sands of Kos
Island indicate elevated rare earth elements (REEs), Ti, and V contents, linked probably back to anandesitic volcanic source, while magnetite separated from metamorphosed small Fe–Ni-laterites
occurrences is REE-depleted compared to large laterite deposits. Although porphyry-Cu deposits
have a common origin in a supra-subduction environment, platinum-group elements (PGEs) have notbeen found in many porphyry-Cu deposits. The trace element content and the presence of abundantmagnetite separates provide valuable evidence for discrimination between porphyry-Cu–Au–Pd–Pt
and those lacking precious metals. Thus, despite the potential re-distribution of trace elements,
including REE and PGE in magnetite-bearing deposits, they may provide valuable evidence for their
origin and exploration.
Keywords: magnetite separates; trace elements; exploration; ophiolites; skarn; porphyry-Cu;
Fe–Ni-laterites

1. Introduction
Magnetite is a widespread mineral, as disseminated or massive ore, in various geotectonic settings
and rock-types, such as magmatic, hydrothermal, ophiolites, calk-alkaline rocks (porphyry-Cu deposits,skarn type), laterites, and metamorphosed rocks/ores, all formed under a variety of conditions [1–13].
It has an inverse cubic (space group Fd3m) spinel-type structure with the general stoichiometry AB2 O4 ,where A and B are tetrahedral (Fe3+ ) and octahedral (Fe3+ and Fe2+ ) coordination sites, respectively [14].
In addition, the geochemistry of magnetite separates is also aﬀected by the co-precipitation of otherminerals [5–9]. Trace elements hosted either in the structure of magnetite, such as cations which cansubstitute on the A (Mg, Fe2+ , Ni, Mn, Co, or Zn) and B (Al, Cr, Fe3+ , V, Mn or Ga) sites [14,15], and/orin mineral inclusions. They may result in distinctive trace element signatures in magnetite separates
and can be used as a prospective tool for many types of ore deposits, and/or a variety of trace-element
discrimination diagrams can classify magnetite separates from diﬀerent types of ore deposits [4–13].
However, it has been proposed that the discrimination diagram (Ti + V) versus (Ca + Al + Mn) can be

Minerals 2019, 9, 759; doi:10.3390/min9120759

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Minerals 2019, 9, 759

used with caution as a petrogenetic indicator for trace element data in the case of magnetite separates
which have unusual composition and/or have been re-equilibrated. [16].
The present study is focused on massive and disseminated magnetite separates from various
geotectonic settings covering a wide range of formation conditions: Hydrothermal magnetite associatedwith calc-alkaline rocks (skarn type, porphyry-Cu–Au deposit), ophiolite complexes (associated withsulphides or apatite), magmatic magnetite from the magmatic sequence of ophiolites and ultramaﬁclavas/wherlites, magnetite from metamorphosed laterites, and magnetite separates from black coastal
sand. Major and minor/trace element contents, including rare earth elements (REEs), in these magnetite
separate types are presented and are compared to literature data. The purpose of this application of themagnetite chemistry and texture features is to provide a potential relationship between compositionalvariation and physico/chemical conditions, the magnetite origin, and exploration of ore deposits.
2. Materials and Methods
2.1. Mineral Chemistry
Polished sections (two-tree samples from each type) from magnetite ores (Tables 1–4) were carbonor gold-coated and examined by a scanning electron microscope (SEM) using energy-dispersive
spectroscopy (EDS). The SEM images and EDS analyses were carried out at the University of Athens(NKUA), using a JEOL JSM 5600 scanning electron microscope (Tokyo, Japan), equipped with the ISIS
300 OXFORD automated energy dispersive analysis system (Oxfordshire, UK) with the following
operating conditions: Accelerating voltage 20 kV, beam current 0.5 nA, time of measurement (deadtime) 50 s, and beam diameter 1–2 μm. The following X-ray lines were used: FeKα, NiKα, CoKα,
CuKα, CrKα, AlKα, TiKα, CaKα, SiKα, MnKα, and MgKα. Standards used were pure metals for theelements Cr, Mn, Ni, Co, Zn, V, and Ti, and Si, MgO for Mg, and Al2 O3 for Al. Contents of Fe2 O3
and FeO for spinels were calculated on the basis of the spinel stoichiometry. The H2 O for (Co, Mn,Ni)-hydroxides was calculated to obtain 2 (OH) pfu.
Whole Rock Analysis
The studied magnetite samples were massive and disseminated mineralizations, derived fromlarge (weighing approximately 2 kg) samples, which were crushed and pulverized in an agate mortar.
The samples of magnetite separates from the Skouries deposit are representative of increasing depth
from 43 to 363 m (the depth for the sample skou8b was 280 m) [4]. The multi-stage water ﬂotation
was applied and magnetite separates were concentrated magnetically from the water suspension of
the rock powder. This portion was furthermore pulverized in the agate mortar at a size fraction −75
μm, and magnetite separates were separated furthermore from silicates. Thus the magnetite separate
concentrates with weak silicates and potential very ﬁne inclusions of various minerals (SEM/EDS
images; Table 1) were used for bulk rock analysis.
Major and minor/trace elements in magnetite separate samples and two samples from the extinct
fumaroles co-existing with native sulphur from the volcanic environment in the Aegean Sea (Santorini
and Nisyros islands) were determined at the SGS Global—Minerals Division Geochemistry Services
Analytical Laboratories Ltd., Vancouver, BC, Canada. The samples were dissolved using sodium
peroxide fusion, combined ICP-AES and ICP-MS (Package GE_ICP91A50). Detection limits of themethod are provided (Table S1). On the basis of the quality control report provided by AnalyticalLabs, the results of the reference material analysis in comparison to expected values, and the resultsfrom multistage analysis of certain samples, showed accuracy and precision of the method in good
agreement with international standard (<10%).

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3. Geological Outline and Mineralogical Features
3.1. Hydrothermal Magnetite Separates
3.1.1. Massive Magnetite Separates of Skarn-Type (Lavrion, Plaka Mine)
The ancient (earlier than 1000 B.C.), famous mine of Lavrion in Attica (Greece) is associated witha granodiorite intrusion, of Upper Miocene age, within the metamorphic -Cycladic Crystalline belt,
which belongs to the carbonate replacement deposits (CRDs) [17]. It is well known for the exploitationof mixed massive sulphide mineralization composed by sphalerite, pyrite and galena (B.P.G), especiallythe production of silver [17–20]. Early hornfels formation was followed by the multistage developmentof a skarn deposit at the area of Plaka, around the granodiorite body, located into two stratigraphic
horizons, near the contacts between the Kaessariani schist with lower and upper marble (Figure 1) [2].

Figure 1. Sketch map showing the location of the studied magnetite separates hosted in calc-alkaline
rocks (a). Ophiolitic rocks and metamorphosed Ni-laterite deposits (b).

Texture characteristics of the magnetite mineralization show multiple stages of the mineralizationand that early magnetite and magnetite–hematite representing endoskarns (stage I) was followedby pyrite, pyrrhotite, and Fe-rich sphalerite, Ag-rich galena, arsenopyrite, forming exoskarns (stageII) [2,16] high-Fe sphalerite, chalcopyrite, and galena forming (Figure 2) [17–19]. Magnetite occurs as
193

Minerals 2019, 9, 759

euhedral or anhedral crystals, homogeneous or having abundant porosity and solid inclusions, mostly
silicate minerals. It is almost pure, having very low Si, Al, and Mn contents, while Ti and V are lower
than the detection limit (Table 1).

Figure 2. Back-scattered electron (BSE) images showing homogeneous magnetite separates (mt) in
sharp contact with magnetite, having abundant porosity and solid inclusions, mostly silicate minerals,
associated with clinopyroxene, quartz, chlorite, pyrrhotite, chalcopyrite, and sphlerite (a); magnetite is
associated with silicates, sphalerite, and pyrite (b). Abbreviations: mt = magnetite; po = pyrrhotite; cp= chalcopyrite; py = pyrite; sph = sphalerite; cpx = clinopyroxene; qtz = quartz; chl = chlorite.

3.1.2. Disseminated Magnetite in Porphyry-Cu Deposit (Skouries)
The Skouries porphyry-Cu deposit is related to tertiary subvolcanic–porphyritic stocks, which
belong to the Vertiskos Formation of the Serbo-Macedonian massif (SMM), Greece, and are controlled
by deep fracture systems [21]. The typical alteration types of the porphyry-Cu intrusions are
more or less presented in the Skouries intrusion, due to the repeated overprinting and intense
siliciﬁcation, with potassic being the predominant alteration type. Two mineral assemblages occurring
as veinlets and disseminations can be distinguished: (1) Magnetite, reaching up to 10 vol.% (average
6 vol.%)—bornite–chalcopyrite, linked to pervasive potassic and propylitic alteration type, in thecentral parts of the deposit; and (2) chalcopyrite–pyrite, dominant at the peripheral parts of thedeposit [4]. Chalcopyrite, and to a lesser extent bornite, contain occasional exsolutions of clausthalite(PbSe) (Figure 3b). There is a frequent association of magnetite separates and Cu-minerals (bornite andchalcopyrite) with inclusions of thorite, U-bearing thorite, rare earth element (REE)-enriched silicatesof the epidote-group (allanite), Ce- and Th-rich monazites, and zircons identiﬁed as accessory minerals(Figure 3b–d; Table 1).

194

100.6

101.5

101

99.6

0.7
n.d.
0.9
0.5
67.6
30.8
n.d.
n.d.
n.d.
n.d.

Total

0.9
n.d
2.1
0.7
66.7
30.6
n.d.
n.d.
n.d.
n.d.

1.7
0.5
n.d.
n.d.
67.5
30.4
n.d.
0.5
n.d.
n.d.

wt.%
SiO2
Al2 O3Cr2 O3
V2 O5
Fe2 O3
FeO
MgO
MnO
CoO
TiO2

2.6
0.9
n.d.
n.d.
67.7
30.4
0.3
0.3
n.d.
n.d.

SG-6

sampleSK43

Skouries

Plaka
P.L.mt1

Plaka
P.L.mt2

LavrionLocation

99.8

0.4
n.d.
n.d.
n.d.
67.9
30.7
0.5
n.d.
n.d.
0.6

Skou8

Hydrothermal

Skarn-type

Deposit type

CALC-ALKALINE

99.3

0.5
0.6
n.d.
n.d.
68.1
30.1
0.9
n.d.
n.d.
n.d.

Agoriani
O.Ag.tOthrys

100.8

2.3
0.2
0.1
n.d.
68.1
29.9
n.d.
0.2
n.d.
n.d.

Samari
EdK2

EdessaHydrothermal

100.6

0.3
n.d.
n.d.
n.d.
69.7
30.2
0.2
0.2
n.d.
n.d.

Perivoli
Pi.Mt

Pindos

99.6

0.5
1.7
n.d.
1.4
55.1
33.5
n.d.
0.5
n.d.
6.9

Krapa
VKMt

Vourinos

OPHIOLITES

99.4

0.9
2.1
n.d.
1.8
55.4
33.3
0.9
0.6
n.d.
4.4

Krapa
VKMt1

Vourinos

Magmatic

100.9

0.3
11.1
3.9
0.7
45.2
32.3
3.1
0.2
0.2
3.6

Agrilia
15Agr2Othrys

100.7

0.7
7.5
1.5
0.6
51.4
31.1
3.5
0.4
n.d.
4.6

Agrilia
O.Agr34

OthrysTable 1. Representative electron SEM/EDS analyses of magnetite in various rock-types (Figures 2–5, 7 and 8). Symbol n.d. = no detected.

100.6

0.5
1.5
0.2
0.5
57.1
34.5
n.d.
1.2
n.d.
5.1

KOS1

KOS

PLACER

Minerals 2019, 9, 759

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Minerals 2019, 9, 759

Table 2. Representative electron SEM/EDS analyses of ﬁne-grained minerals associated with magnetite(Figure 3) from the Skouries porphyry-Cu deposit. Symbol n.d. = no detected.
wt.%

Monazite Thorite

U-Thorite

Na, Ce-epidote

Allanite

Allanite

SiO2
Al2 O3
FeO
CaO
P2 O5
ZrO2
La2 O3Ce2 O3
Nd2 O3Pr2 O3
ThO2
UO3

1.1
n.d.
n.d.
0.8
30.3
n.d.
17.2
31.6
10.5
2.8
5.1
n.d.

18.8
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
79.4
n.d.

15.7
n.d.
0.6
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
70.8
11.7

38.3
27.2
8.3
22.2
n.d.
n.d.
0.9
2.2
n.d.
n.d.
n.d.
n.d.

32.1
16.2
14.9
12.4
n.d.
n.d.
5.1
14.5
4.7
n.d.
n.d.
n.d.

35.4
14.1
17.4
13.6
n.d.
n.d.
3.9
12.5
2.2
n.d.
n.d.
n.d.

Total

99.4

98.2

98.1

99.1

99.9

99.1

Figure 3. Back-scattered electron (BSE) images from the Skouries porphyry deposit, showing
textural relationships between co-existing minerals. Quartz–chalcopyrite vein crosscutting biotite
and magnetite-rich potassic altered porphyry (a); a close intergrowth of disseminated Ti–magnetite,chalcopyrite, bornite, and clausthalite (PbSe) (b); magnetite associated with thorite and zircon hosted byquartz (c); rare earth element (REE)-enriched epidote overgrown by allanite (d). Their composition is
given (Table 2). Abbreviations: qtz = quartz; ab = albite; bi = biotite; mt = magnetite; cp = chalcopyrite;bn = bornite; PbSe = clausthalite; zr = zircon; ep = epidote; all = allanite.

3.1.3. Massive Magnetite Associated with Fe–Cu–Ni–Co Sulphides in Ophiolites (Pindos)
Sulﬁde mineralization in the Pindos ophiolite complex is located in the Smolicas Mountains, near
the Aspropotamos dismembered ophiolite unit. This sequence, belonging to the ophiolitic unit of thePindos complex, includes: (1) An intrusive section composed of dunites, lherzolites, olivine-websterites,
olivine-gabbros, anorthosite gabbros, gabbros, and occasionally gabbronorites; and (2) a volcanic and196

Minerals 2019, 9, 759

subvolcanic sequence composed mainly by basalts and basaltic andesite pillow lavas ranging fromhigh- to low-Ti aﬃnity [22,23]. At the Perivoli (Tsoumes Hill) of the Aspropotamos unit, within gabbro
close to its tectonic contact with serpentinized harzburgite, small irregular to lens-like occurrences (4 m
× 1.5 m) of massive sulﬁde mineralization of Fe–Cu–Ni–Co type have been described [24]. Magnetite,
forming often a network texture, is associated with sulphides, either as massive ore with inclusions of
sulﬁdes (chalcopyrite, pyrite, and pyrrhotite), or as individual grains dispersed within sulﬁde ore. A
salient feature of magnetite is a network texture with occasionally deformed crystals (Figure 4). The
composition of magnetite separates is characterized by very low Al, Ti, Cr, Mn, Ni, and V contents(Table 1 and Table S1).

Figure 4. Back-scattered electron (BSE) images showing elongated curved magnetite associated withsulphides (a,b) from the Perivoli area, Pindos ophiolite complex. Abbreviations as in Figure 3.

3.1.4. Massive Magnetite in Maﬁc Ophiolitic Rocks Edessa (Samari)
The isolated ophiolite masses of Eastern Edessa belong to the Axios zone (Almopias
subzone) [25–27]. At the area of Samari, Edessa, with dominant rock types diabase, gabbros, andserpentinized peridotites, ore bodies consisting of magnetite separates and sulphides have been
described [27]. Magnetite occurs (1) as disseminated euhedral and/or subhedral crystals, along withsulphides (pyrite, chalcopyrite) within gabbros and diabase, accompanied by epidote, chlorite, quartz,
sphene and tremolite; and (2) as massive ore hosted within gabbros, close to its contact with peridotite,with a gangue silicate mass dominant by greenalite, while sulphide veinlet crosscutting the magnetite
separate ore are common (Figure 5) [27]. The magnetite is pure, characterized by the presence of only
negligible contents of other elements (Table 1).

Figure 5. Back-scattered electron (BSE) images showing massive magnetite associated with little
amounts of pyrite in a matrix of chlorite (greenalite) (a) and inclusions of chlorite in magnetite (b) fromthe Samari area (E. Edessa). Abbreviations as in Figures 2 and 3.
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3.1.5. Massive Magnetite Associated with Apatite in Ophiolites
Major Mesozoic ophiolite complexes in Greece, including the Othris, Pindos, and Vourinos suites,
were all obducted from the Middle-to-Late Jurassic and have been interpreted as parts of the same
oceanic slab, although fore-arc and back-arc ridges have been tectonically fragmented and separated
by overlapping sediments [28,29]. The Othrys ophiolite complex consists of a stack of thrust sheets inwhich overlapping stratigraphic successions, consisting mainly of harzburgites, lherzolites, gabbros,
maﬁc dikes, and pillow lavas, have been recognized, while small irregularly shaped bodies of gabbroand gabbroic dikes or veins intrude the plagioclase lherzolite of the mantle sequence [30]. Irregular tolens-like occurrences (0.5 m × 1 m) of massive pure magnetite associated with large (up to 3 cm long)well-formed crystals of chlor–hydroxyapatite apatite, and lesser amounts of serpentine, tremolite, andNi silicates (nepouite, pimelite), and by Ni sulﬁdes (pentlandite, violarite, heazlewoodite) have been
described in the Agoriani area of the Othrys ophiolite complex, central Greece [31]. The analysis ofmagnetite separates showed very little Al, Mg, Ni, Ti, and REE content (Table S1).
3.2. Magmatic Magnetite Separates
3.2.1. Disseminated Magnetite in Ultramaﬁc Lavas and Wherlites
The Agrilia formation, Othrys ophiolite complex, about 6 km Northwestern of Lamia, is an unusual
ultramaﬁc lava characterized by high Mg content (31–33 wt.% MgO) [32]. Olivines from the Agrilia
ultramaﬁc lavas display high forsterite (Fo) contents, and calculated parental magma (in equilibriumwith Fo90.5 ) approximately 17 wt.% MgO [33]. Disseminated chromite occurs within groundmass(devitriﬁed glass) and as inclusions in olivine and clinopyroxene [32–35]. The chromite spinel within
groundmass is commonly rimmed by Ti–magnetite with sharp contacts (Figure 6a). Chromite is
Cr-rich, exhibiting a limited range of variation (Table 3). Amphibole (tschermakite after [34], rhönite,
chromite, and Ti–magnetite all associated with groundmass (Figure 6)). Rhönite in association with
Ti–magnetite within devitriﬁed glass (Table 3), and inter-grown with amphibole, may be of genetic
signiﬁcance [33,34]. The above ultramaﬁc lavas are associated with medium-to-coarse grained wehrlite.
They are mainly composed by olivine, clinopyroxene, hornblende, and in lesser amounts, spinel,
orthopyroxene, phlogopite, and magnetites.
Table 3. Representative electron SEM/EDS analyses of rhönite (Rho) and amphibole (amph) within
devitriﬁed glass (Figure 6) from the Agrilia ultramaﬁc lavas [35]. Symbol n.d. = no detected.
Host

Devitriﬁed Glass (d.G.)

SiO2
MgO
Al2 O3Cr2 O3
CaO
TiO2
FeO
MnO
Na2 O
K2 O
NiO
Cl

Rhonite (Rho)
26.3
26.6
8.8
8.7
15.7
14.9
n.d.
n.d.
10.7
10.6
4.9
3.9
28.9
29.8
0.2
0.2
0.5
0.5
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.

Total

96.0

95.2

25.8
8.8
15.2
n.d.
10.6
4.7
30.3
0.3
1.3
n.d.
n.d.
n.d.

d.G.
33.1
14.5
12.1
0.8
4.3
1.6
27.1
0.2
1.6
0.2
n.d.
n.d.

45.6
6.5
9.9
1.1
3.1
0.2
21.8
0.2
6.9
0.9
n.d.
n.d.

39.8
9.2
16.5
n.d.
10.1
1.9
17.7
n.d.
2.5
0.3
n.d.
0.2

Amphibole (amph)
40.4
10.6
17.5
n.d.
9.6
1.2
17.1
0.3
1.7
0.2
n.d.
0.2

55.2
11.1
10.6
1.3
10.4
0.2
6.2
n.d.
2.7
0.3
n.d.
n.d.

97.0

95.5

96.2

98.2

98.8

98.0

3.2.2. Disseminated Magnetite Separates in Norite Gabbros (Central Vourinos, Krapa Hills)
The Vourinos ophiolite complex located in NW Greece constitutes a complete but tectonicallydisrupted ophiolite sequence. Two magmatic series have been preserved at the central part of thecomplex, which constitute the earlier Krapa sequence, consisting of magnetite-bearing gabbro-norite
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Minerals 2019, 9, 759

and leuco-gabbro-norite (minor magnetite-bearing) and the younger Asprokambo sequence [36]. The
Fe–Ti mineralization occurs as disseminated grains in a proportion of 5–20 vol.% between coarse-grained
clinopyroxene and plagioclase, while sulphides (pyrrhotite, pyrite, chalcopyrite) are present in lesser
amounts. Ti–magnetite occurs as interstitial to the major phases as small disseminated euhedral
and irregular grains (Figure 7). Ti-bearing magnetite separates exhibits low Si, Mg, and high Al(1.7–2.3 wt.% Al2 O3 ), Ti (3.5–6.9 wt.% TiO2 ), and V (1.4–2.1 wt.% V2 O5 contents) (Table 1).

Figure 6. Back-scattered electron (BSE) images from ultramaﬁc lavas of the Agrilia Formation, showing
the association of Ti–magnetite (Ti–Mt) with chromite (chr) (a), microcrystals of rhönite (Ron) (b)associated, and randomly oriented amphibole (Amp) (c,d).

3.3. Magnetite in Coastal Black Sand
The volcanic rocks of the South Aegean arc form a chain from the Saronikos Gulf in the west,
to Kos, Nisyros, and the east, through the Milos and Santorini islands [37]. Multiple eruptions ofcalc-alkaline to tholeiitic composition and Plio-Pleistocene age are located from the Santorini toward
the Nisyros–Kos volcanic rocks (Figure 1a) [37–39]. Those volcanic rocks are part of the volcanic
system related to the northward subduction of the last remnant of the oceanic crust of the African plate
beneath the southern edge of the active margin of the European plate [38–41]. Magnetite and otherheavy minerals have been found in all volcanic types in small quantities (1–2 modal %).
The investigated coastal black sand samples (black sea sand) collected from the Northern partof Kos Island revealed the presence of abundant Ti–magnetite with inclusions of apatite, ilmenites,
Fe–Mn-oxides and monazite [(Ce, La, Th, Nd, Y) PO4 ] (Figure 8). Feldspars and quartz are common too.

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Minerals 2019, 9, 759

Figure 7. Back-scattered electron (BSE) images from the gabbro-norite sequence of Central Vourinos(Krapa area) showing disseminated Ti–magnetite as small euhedral and irregular phase associatedwith ilmenite (il) interstitially to the major phases of clinopyroxene (cpx) anf plagioclase (pl) (a–d).

Figure 8. Back-scattered electron (BSE) images of coastal black sand collected from the northern part of
Kos Island, showing that Ti–magnetite is the dominant oxide, while apatite (apat) forms inclusions ofapproximately 100-μm-long grains (a–c); Fe–Mn-oxides (b), monazite [(Ce, La, Th, Nd, Y) PO4 and (Ce,La, Nd, Pr, Sm) PO4 ] are common (a,d).
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Minerals 2019, 9, 759

The sample from the Nisyros caldera (neighboring Kos Island) comes from whitish extinct
fumaroles co-existing with native sulphur, that is dominant by opal (Figure 9). The analyzed sampleof the formation from the Nisyros caldera is opal–CT and quartz [41], which it is associated with
abundant native sulphur (Figure 9).

Figure 9. Back-scattered electron (BSE) images (unpolished sections) of extinct fumarole formationfrom the hydrothermal crater of the Nisyros caldera, showing native sulphur (a,b) and the associationof native sulphur with whitish Si material (opal) (c), having low REE content (Table S1).

3.4. Magnetite Separates from Metamorphosed Fe–Ni-Laterites (Olympos, Edessa, and Vermion)
The dismembered ophiolite masses of the Olympos, Edessa, and Vermio mountains (Upper
Jurassic–Lower Cretaceous), composed mainly by serpentinized harzburgite and, to a lesser extentm
crustal magmatic rocks (pyroxenites and gabbros), outcrop along the eastern margin of the Pelagonian
massif. They are considered to have been derived from the Almopias subzone of the Vardar zone
and were overthrust westward onto the Pelagonian massif [22]. They are mostly composed of amantle sequence (harzburgites) and crustal magmatic rocks, both characterized by supra-subduction
zone (SSZ) features [23]. Small laterite bodies are located in the Olympos, Edessa, and E. Vermio
Mountains, at the contact of serpentinized harzburgites with Upper Cretaceous limestones or within
the serpentinites themselves. Due to intense tectonism, the Fe–Ni-laterite occurrences are often entirely
enclosed within serpentinized harzburgites [1,25]. Apart from ﬁne-grained magnetite separates in
the matrix of the Fe–Ni ore, magnetite separates are associated with chromite, as the product of a
gradual alteration of chromite or epitaxial as rims on chromite (Figure 10). A salient feature of zoned
chromite grains from the Olympos, Edessa, and Vermion laterite occurrences is a Cr, Al, and Mg
gradual decrease as Fe increases outward from the core; Mn, Zn, and Co increases gradually outwards,
attaining greatest values at the periphery of chromite cores and in the Fe–chromite, up to 13.0, 2.1,
and 4.1 wt.%, respectively, and drops oﬀ to negligible values in magnetite separates (Table 4). At
the area of Vermio, there is abundant garnet (grossularite) and calcite, while Ni is mainly hosted inchlorite, serpentine, and theophrastite containing 80 wt.% NiO (Figure 10; Table 4) [42]. A (Co, Mn,Ni)-hydroxide with a wide compositional range occurs in a spatial association with theophrastite(Figure 11; Table 4) [43]. Several SEM-EDS analyses of hydroxides can be classiﬁed into two groups(Figure 11): Group A includes phase mineral dominated by cobalt, and group B those dominated by
Mn, towards the peripheral parts of the concentrated development (Figure 10).

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Minerals 2019, 9, 759

Figure 10. Back-scattered electron (BSE) images of metamorphosed Ni-laterites from northern Greece.
Magnetite (mt) surrounding chromite (chr) grains and occurring along cracks (a,b,e,f); various
morphological forms and textural relationship between magnetite–theophrastite (Th) and garnet (grt)(c,f,g); magnetite crosscutting veinlets (e,i,j); magnetite–(Co–Mn–Ni)(OH)ss showing successive thinlayers, composed by ﬁne ﬁbrous crystals (d); microtextures resembling fossilized microorganisms
coated by (Co–Mn–Ni)(OH)ss (white arrows) (f) or isolated and aggregates of bacterio-morphic
Fe–Ni–serpentine (white arrows) within a veinlet of Ni–serpentine (Ni–srp) (g); chromite–Fe–chromite(h); chromite–magnetite (i); magnetite–apatite (ap) (j).

202

Mineralsamplewt.%
SiO2
Al2 O3Cr2 O3Fe2 O3
FeO
MgO
MnO
ZnO
CoO
TiO2
NiO
Total
Cr/(Cr + Al)
Mg/(Mg + Fe2+ )Location

Fe-chr
L1

n.d.
8.6
53.7
3.5
27.6
0.6
4.1
0.4
n.d.
n.d.
n.d.
98.5
0.81
0.03

Core
Chromite
L1

0.2
8.2
62.1
1.8
13.9
11.3
1.9
n.d
n.d.
n.d.
n.d
99.4
0.83
0.55

Paliambela

Olympos

0.4
0.4
5.9
62.2
28.4
0.4
0.5
n.d.
0.4
n.d.
1.6
100.2
0.81
0.13

Rim
mt
L1
n.d.
9.7
60.2
2.7
15.5
11.2
1.6
n.d.
n.d.
n.d.
n.d.
100.9
0.78
0.47

Core
Chromite
Ed10
n.d.
6.5
58.1
3.3
8.3
3.5
13.1
2.1
4.1
n.d.
n.d.
99.0
0.86
0.17

Fe-chr
Ed10
0.5
0.2
1.6
67.7
29.3
0.3
n.d.
n.d.
n.d.
n.d.
1.4
101
0.87
0.01

Nissi
Rim
mt
EdP10

Edessa

0.3
24.4
43.3
2.2
17.3
11.1
n.d.
0.5
n.d.
n.d.
n.d.
99.1
0.54
0.51

Core
Chromite
V.AK
0.7
7.7
42.3
17.6
26.6
2.2
2.8
0.7
n.d.
n,d.
0.2
100.8
0.79
0.12

Fe-chrV.AK

Mavrolivado

1.2
0.5
1.6
66.5
28.9
0.3
0.3
0.3
n.d.
n.d.
n.d.
99.6
0.94
0.01

Rim
mt
V.AK

E.
Vermion

n.d.
n.d.
n.d.
n.d.
1.0
n.d.
n.d.
n.d.
n.d.
n.d.
79.2
80.2

n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
25.0
n.d.
38.0
n.d.
17.4
80.4

n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
13.9
n.d.
54.3
n.d.
11.2
79.44

Theophastite(Co,Mn,Ni)-hydroxides

Stournari

n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
19.2
n.d.
46.9
n.d.
13.6
79.7

Table 4. Representative electron SEM/EDS analyses of chromite cores, magnetite (mt) and Fe–chromite (Fe–chr), Theophrastite and (Co, Mn, Ni)-hydroxides from
Ni-laterites of the Vermion mountain (Northern Greece; Figure 10). Symbol n.d. = no detected.

Minerals 2019, 9, 759

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Minerals 2019, 9, 759

Figure 11. Triangular plot of Co, Mn, and Ni (atom%) showing the compositional variation of the (Co,Mn, Ni)solid solutions [43].

The (Co, Mn, Ni)solid-solution phases are associated with theophrastite, pyrochroite, Ni–serpentine(garnierite), garnet (grossularite), and magnetite, all having a common origin, subsequently of thestrong late tectonic evolution, overprinting earlier deformation events (Figure 10c–g). They occur
as successive thin (a few to tens of μm layers, composed by ﬁne ﬁbrous crystals up to a few μm(Figure 10d), microtextures resembling fossilized microorganisms coated by (Co–Mn–Ni) (OH)ss (whitearrows) (Figure 10f), or isolated and aggregates of bacterio-morphic Fe–Ni–serpentine (white arrows)
within a veinlet of Ni–serpentine (Ni–srp) (Figure 10g).
4. Geochemical Characteristics of Magnetite Separates
Although a good separation of magnetite from very ﬁne inter-grown silicates and/or sulphides(Figures 2–10) was not achieved, it is clear that the highest Cr, Mn, and Co contents were recorded
in Fe–Ni-laterite occurrences of northern Greece, ranging from 1.1 wt.% Cr in Vermio to 9.0 wt.% Cr
in Olympos, from 0.38 wt.% Mn in Vermio to 0.69 wt.% Mn in Olympos, and from 460 ppm Co in
Olympos to 1060 ppm Co in the Edessa laterites (Table S1).
The highest Ti and V contents were determined in disseminated magnetite separates from themagmatic sequence (norite gabbros) of the Vourinos (Krapa) ophiolite complex (2.56 wt.% Ti, 6330
ppm V), the Agrilia ultramaﬁc rocks from the Othrys complex (7.48 wt.% Ti, 1760 ppm V), and thecoastal black sands (or placer deposits) from Kos Island. In addition, magnetite separates from theAgrilia ultramaﬁc rocks is richer in Al, Mg, Ti, and Cr compared to that from the Krapa norite-gabbros.
Magnetite separates from the Skouries deposit shows relatively signiﬁcant contents of Ni (up to 640
ppm Ni) and Cr (up to 1060 ppm Cr) (Table S1).
Hydrothermal magnetite separates from the ancient Lavrion (Plaka) mine, that is of skarn-type,
and the Skouries porphyry deposits, that is disseminated, both associated with calc-alkaline rocks
diﬀer in terms of the higher Ba, Bi, Cr, Ni, Co, V, Sr, Sc, Rb, ΣREE, Y, Hf, Th, U, and Zr contents
in the latter (Table S1). Gallium is higher in magnetite separates associated with cal-alkaline rocks(38–59 ppm) compared to that associated with ophiolites (2–21 ppm Ga) (Table S1). The highest Ga
contents were recorded in magnetite separates from the Skouries porphyry-Cu deposit (40–59 ppm),
skarn-type from Lavrion mine (38–39 ppm), and coastal sand from Kos Island (42 ppm), all related with
calc-alkaline rocks. A potential order of the average Ga content in magnetite separates, from low to highis from Ni-laterites, hydrothermal (ophiolites) associated with sulphides, ultramaﬁc lavas (ophiolites),hydrothermal (ophiolites) associated with apatite, magmatic (norite-gabbros), skarn-type (calc-alkaline),
coastal sands (calc-alkaline volcanic), and porphyry-Cu deposits (hydrothermal calc-alkaline), resulting
into a discrimination using a plot of (V + Ti) versus Ga content (Figure 12).
The rare earth element (REE) content of the majority of hydrothermal magnetite separates analyzed,
such as those from the Lavrion mine (skarn-type), the Skouries porphyry-Cu deposit, the Pindos
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Minerals 2019, 9, 759(Perivoli) and Edessa (Samari) is generally low (a few decades ppm and REEs exhibit similar trends.
Although the REE content, as well as actinides (Th, U) and Zr, in the porphyry (Skouries) deposit are
higher compared to those in skarn-type (Lavrion), the REE patterns are similar in terms of the negative
Eu, Tm, and Ho anomalies, but the former display occasionally positive Eu anomalies as well, andslight negative anomalies in Sm and Tm (Figure 13a,b). Moreover, considering the magnetite separate
samples from the Skouries deposit, it seems likely that there is a small increase of the REEtotal with
the increasing depth from 43 to 363 m. In addition, the magnetite separate sample from the lowest
depth (43 m) exhibits a positive Eu anomaly, in contrast to the samples from deeper parts of the deposit(Figure 13; Table S1). The hydrothermal magnetite separates associated with apatite from the Othrys(Agoriani) complex show higher REE content (54 and 80 ppm), that is, increasing with the presence of
apatite, reaching values to more than 1300 ppm REE in apatite [31].
Calculations of Ce, Eu, Pr, and Gd anomalies were deﬁned [44] by equations such as: Ce/Ce*= (2Ce/Ce chondrite)/(La/La chondrite + Pr/Pr chondrite) and Eu/Eu* = (2Eu/Eu chondrite)/(Sm/Sm
chondrite + Gd/Gd chondrite). Assuming that values > 1 and < 1 are called positive and negative
anomalies, respectively, chondrite normalized REE patterns for Ni-laterites from the Vermio mountains
showed negative Ce, Eu anomalies, which is in agreement with the Ce/Ce* (0.56–0.24) and Eu/Eu*(0.40–0.24) values (Table S1). The highest REE contents were recorded in Ti–magnetite separates fromthe ultramaﬁc lavas (130 ppm) and from the coastal sands, Kos Island (220 ppm). Sulphur occurs to
be impregnated with particles of rocks, including opal–CT, quartz, alunite, anhydrite, and kaolinite(Figure 9) [41]. There is enrichment for light rare earth elements (LREE) from La to Sm and decrease
for the heavier REE in magnetite separates compared to andesitic volcanic rock from Santorini Island(Figure 13f), and enrichment for all REE compared to the recent formation, associated with abundantnative sulphur in Nisyros Island (Figure 9a,b). A distinct negative Eu anomaly is a common feature of
the magnetite separates from the coastal sand of Kos and the andesitic rock from Santorini Island.

Figure 12. Cont.

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Minerals 2019, 9, 759

Figure 12. Plots of (V + Ti) and Al versus Ga contents (a,b), Eu/Eu* ratio versus Ga content (c); (Ce+ Eu) versus Zr content (d), Eu/Eu* versus Gd/Gd* ratios (e), all for magnetite separates (data from
Table S1) and plot of the Eu/Eu* ratio versus Ga content for Fe–Ni-laterites from Northern and Central
Greece (Ni-lat), bauxite latererite (B-lat), and bauxites from Parnassos-Giona zone (f). Data from Table
S1 and literature [45,46].

Figure 13. Chondrite normalized diagrams for magnetite separates (a–f) and volcanic rock (d,f). Data
from Table S1 and [47] for ultramaﬁc lavas (sample labeled as Rock UL, n = 3, and wehrlite (Wh), n = 7)
from the Othrys ophiolite complex. Chondrite values from McDonough and Sun [48].

Moreover, the same trace elements were analyzed in two samples representing extinct fumarolesof andesitic and rhyolite composition, co-existing with native sulphur from the Aegean active volcanic

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Minerals 2019, 9, 759

arc, collected from Santorini (andesite) and Nisyros Island (Figures 9 and 13f). There is enrichment of
LREE from La to Sm and decrease of the heavier REE in magnetite separates from costal sands (Kos),that is comparable to andesitic material from Santorini Island (Figure 13f). In addition, they exhibit
elevated REE content compared to the extinct fumarole from Nisyros Island (Table S1; Figure 12f).
Furthermore, REE patterns of both extinct fumarole material and native sulphur from Nisyros diﬀer interms of the low LREE content, the relative HREE enrichment compared to magnetite separates andandesitic material from Santorini and the negative Ce, Pr, and Eu anomalies (Figure 13f; Table S1).
Relatively high REE contents were recorded in Ti–magnetite separates from the ultramaﬁc lavas
from the Agrilia (99 ppm), while in wehrlite it is only 23 ppm. However, both are clearly higher
compared to ultramaﬁc lavas and wehrlite rocks (literature data), ranging from 14 to 17 ppm REEtotal ,respectively [47]. Chondrite normalized REE for magnetite separates from ultramaﬁc lavas andwehrlite display sub-parallel patterns, with slight LREE enrichment in ultramaﬁc lava, and an inverse
relationship in terms of heavy rare earth elements (HREE) (Figure 13d). The Eu content in Ti–magnetite
separates from the ultramaﬁc lavas of Agrilia (Othrys) is enriched, showing a positive anomaly; incontrast, it is depleted in the magmatic sequence (norite gabbros) of the Vourinos (Krapa) complex(Figure 13d). In addition, the REE patterns for ultramaﬁc lavas and wehrlites (literature data [47]) show
a slight HREE enrichment for wehrlites compared to ultramaﬁc lavas too (Figure 13d), while bothrock types exhibit clear positive Ce, Nd, Gd, and Er anomalies and negative Pr, Eu, and Tb anomalies(Figure 13d).
In summary, the analyzed samples of magnetite, both massive ore and separates, covering a wide
range of associated rock type, reveal wide compositional variations (Table S1 and Table 5). Althoughmuch compositional overlaps were recorded, a potential classiﬁcation of magnetite separates is: (a)Magmatic: (a1) Associated with ophiolites and (a2) associated with calc-alkaline volcanic rocks (coastal
sands); (b) hydrothermal: (b1) Associated with ophiolites and (b2) associated with calc-alkaline rocks(skarn-type and porphyry-Cu–Au deposits); and (c) epigenetic origin of magnetite separates in laterites(Figures 2–13). Average values show the highest ΣREE contents in disseminated magnetite separates
from ultramaﬁc lavas (Othrys) and coastal black sands (Kos Island) (Table 5).
Available data from previous publications have shown that an enrichment in Cr and Ni is acommon feature of porphyry-Cu–Au–Pd–Pt, such as the Skouries, Elatsite, and other porphyry deposits
worldwide, characterized by signiﬁcant Pd and Pt contents, in contrast to the porphyry-Cu ± Mo
deposits of Russia and Mongolia with lower Pd–Pt content (Figure 14) [4,12–14,49].

207

Fe-Ni-laterite
Bauxite laterite
BauxiteMETAMORPHOSED

Calk-alkaline

HYDROTHERMAL
Ophiolites

Calc-alkaline

Rock-Type

TextureMajor, matrix
Rare, matrix
Rare, matrix
Rare, matrix

Ni-laterites (n = 4)

Ni-laterites (n = 13)
Bauxitic laterites (n = 6)
Bauxites (n = 17)

Olympos,Vermio,
EdessaLokrisLokrisParnassos-Giona

Massive
Massive
Massive
DisseminatedMassive

Disseminated
DisseminatedPlacer

Peridotite
Association with apatite
Irregular
Porphyry-Cu
Skarn

Norite gabbro
Ultramaﬁc lava
Coastal sandEdessa (Samari)Othrys (Agoriani)
Pindos (Perivoli)
Skouries
Lavrion (Plaka)

Vourinos (Krapa)
OthrysKos

Ophiolites

120
350
470

37

10
76
3.1
26
3

6
132
224

ΣREE

1.02
1.01
2.64

0.41

0.94
0.82
0.90
0.98
0.81

1.06
0.9
0.92

Ce/Ce*

0.85
0.71
0.69

0.32

0.11
0.70
0.74
1.01
0.48

1.43
1.97
0.32

Eu/Eu*

0.96
0.97
0.58

1.45

0.95
0.71
1.02
0.88
0.94

0.92
0.98
0.85

Pr/Pr*

1.14
1.2
0.8

0.93

1.75
1.28
1.34
1.20
1.64

1.14
0.65
1.81

Gd/Gd*

7.6
17
70

4.2

2
20
9
46
3.9

22
12
42

Ga

75
340
460

280

34
640
80
760
44

6520
1820
1480

V

2900
7200
16,000

380

0.02
0.02
<0.01
0.25
0.02

2.6
13.4
2.5

Ti

1900
3500
820

33,980

560
1100
12
630
160

1070
1600
1150

Cr

Table 5. Summary of the certain trace element contents (ICP-MS analyses) in magnetite separates and/or host ores.

Location

MAGMATIC

Ni

2200
1400
150

9800

56
280
6
340
30

100
160
11

Reference

[46]
[46]
[45]

Present study

Present study
Present study
Present study
Present study
Present study

Present study
Present study
Present study

Minerals 2019, 9, 759

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Minerals 2019, 9, 759

Figure 14. Plot of the (Pt + Pd) versus Cr content in porphyry-Cu–Au–Pd–Pt and porphyry-Cu–Mo
deposits. Data from literature [12,14,49].

5. Discussion
The trace element content in magnetite separates, coupled with mineral chemistry and texture
relationships, may provide evidence on (1) their parent magma or hydrothermal ore forming system,(2) partition coeﬃcient of trace elements into magnetite separates, and (3) physico/chemical conditions(temperature, fO2 ) during deposition and/or re-deposition of magnetite separates, and can be used as aprospective tool for many types of ore deposits [5–10,50]. The recorded wide variations in the traceelement contents of massive and concentrates of disseminated magnetite separates from Greece may
be used to discriminate between magmatic and hydrothermal magnetite separates, associated eitherwith calc-alkaline rocks and ophiolites, between porphyry-Cu + platinum-group elements (PGEs) andporphyry-Cu lacking precious metals deposits, and between laterite types.
5.1. Trace Element Distribution in Magnetite Structure and Associated Minerals
The geochemistry of magnetite separates is aﬀected by a variety of cations accommodated in thespinel-type structure [3,14,15] and the co-precipitation of other minerals [5–10]. The high-resolution
studies of minerals applying TEM techniques provide a new and very important way to elucidate
valuable information on the physico/chemical processes and mechanisms during formation of ore
deposits and post-depositional changes of ore minerals [51]. The investigation of hydrothermal
magnetite separates from the Los Colorados iron oxide–apatite deposit in Chile showed three
generations of magnetite, which vary signiﬁcantly in terms of their Si, Al, Ca, Mg, Ti, Mn, Na,
and K minor and trace element contents, in contrast to the V content (remarkably constant), andconcluded a complex incorporation of trace elements during growth of magnetite [51]. Although,there is a discrepancy between experimentally determined values and those derived from natural
occurrences, the partition coeﬃcients for spinel–liquid pairs (Dsp/lq ), for high ﬁeld strength elements,
such as Ga, V, Zr, Zn, and Co, are extremely low [50,52]. Considering the presented data (Table
S1), potential components of sulphide (Cu, Zn) and/or silicate (Si, Mg, K) inclusions in magnetite
separates are used with caution. Selected elements, such as V, Ti, and Ga, in individual types ofmagnetite separates are plotted, showing diﬀerent ﬁelds for magmatic and hydrothermal magnetite,and furthermore, between hydrothermal magnetite associated with ophiolites and calc-alkaline rocks(Figure 12).

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Hydrothermal apatite from the Othrys (Agoriani) ophiolite complex contain over 1300 ppm
ΣREE [31], whereas the associated magnetite separates contain less than 100 ppm ΣREE, due probably
to incomplete separation of apatite (low P and Ca contents) (Table S1) [31]. Similarly, magnetite
separates from the coastal black sand from Kos, containing inclusions of apatite (Figure 8b,c), exhibit
relatively high (over 220 ppm ΣREE) (Table S1) and monazite of varying composition [(Ce, La, Th,
Nd, Y) PO4 and (Ce, La, Nd, Pr, Sm)PO4 ] (Figure 8a,d). Magnetite separates from black coastal sands
can be linked back to a volcanic source, and may be a potential REE source [41]. On a global scale,in many placer deposits, such as the river placers in Malaysia, paleo-placers in South Africa, Elliot
Lake, Ontario, and beach sand deposits in India, monazite is an important REE mineral and many
monazite ore deposits are economically important REE resources, although small contents of Th andU make monazite a restricted mineral in India [53]. Also, hydroxyl–bastnaesites and a (La, Nd, Y)(CO3 ) F member of the bastnaesite group, which are among the most economically important as well,have been identiﬁed at the lowest part of karst-type laterites in Greece, close to their contact withcarbonates [54–57].
The ﬁrst data on magnetite separates from black coastal sand from Kos Island are comparedwith potential parental rocks of the neighboring active Nisyros Island and Santorini, all belonging
to the Aegean volcanic arc, showing REE, Ti, V, Cr, Mn, Zn, and Ga contents in magnetite separates
are higher than those in the volcanic rocks of basaltic composition from Santorini Island [50] andNisyros Island (Table S1). The similarity between the REE patterns, including the clear Eu negative
anomaly of magnetite separates from coastal black sand of Kos Island and that of andesitic material
from Santorini Island (Figure 13f), may provide evidence of an andesitic source of the placer magnetite
separates in Kos. Also, the composition of native sulfur has been used to qualitatively estimate
elemental abundances in low-temperature fumarole gases. At low temperatures (<250 ◦ C), oxygen in
the atmosphere interacts with H2 S, resulting to precipitation of native sulphur [58].
5.2. Re-Mobilization and Re-Deposition of Magnetite
Magnetite separates as a late magmatic or early hydrothermal phase in both calc-alkaline systems
and ophiolite complexes may be dissolved and re-precipitated, due to changes in temperature,salinity, pressure, and f O2 [51,59–61]. An increase in temperature would enhance the solubility and,
consequently, the undersaturation of iron in the ﬂuids [61], leading to dissolution of the primarymagnetite separates and formation of the secondary varieties. Thus, primary magnetite may have been
re-mobilized during late brittle deformation and motion along the detachments contacts [60]. In general,
textural and compositional data for magnetite from nine iron skarn deposits in Canada, Romania, andChina have shown that most samples have been re-equilibrated by dissolution and re-precipitation
and/or recrystallization [16,51]. These authors concluded that the applicability and reliability of traceelement discrimination diagrams may be used with caution where magnetite has been subjected tomultiple re-equilibration processes, and that detailed textural characterization must be undertaken inorder to evaluate the eﬀects of hydrothermal overprinting events and other re-equilibration processeson the trace element compositions of magnetite.
A peculiarity of the above Fe–Ni-laterite occurrences is the relatively high Cr, Mn, Co, and Ni
content in Fe–chromite zones of spinels. The epigenetic mobilization of Fe, Mn, Co, and Ni and theirre-deposition is suggested by a signiﬁcant Mn, Zn, and Co enrichment in Fe–chromite (Table 1) andthe abundance of Ni–serpentine (Figure 10; Table 4). The substitution for Mg2+ and Fe2+ by Mn, Zn,
and Co in the chromite lattice resulted to a gradual transformation of chromite to Fe–chromite, whilein the magnetite, that is considered to be the ﬁnal product of this alteration process [1,62] and theabove elements drop to negligible amounts (Table 4). Calculations for spinel structures [14,15,63,64]
have shown that in normal spinels, there is a tendency for the electrostatic energy to favor the larger
ions (Fe2+ > Mg2+ and Fe3+ > Al3+ ) into the tetrahedral site. Thus, the resulting small expansionof the tetrahedral and octahedral sites hampers the entry of other trace elements in the magnetite
separate lattice at low temperatures [65,66]. Besides, the re-mobilization of Fe and re-deposition

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Minerals 2019, 9, 759

of magnetite at low temperature is indicated by its association with (Ni, Fe, Co, Mn)-hydroxides.
Speciﬁcally, the texture relationships between magnetite, theophastite [Ni(OH)2 ], and (Co, Fe, Mn, Ni)(OH)2 (Figure 10a–h) suggest a common mineral forming system, along shear zones, subsequentlyof the strong late tectonic evolution, which overprints earlier deformation events, probably at low
temperature, approximately 100 ◦ C [67] and alkaline conditions, as it is suggested by the presenceof calcite and the high Ca content in the magnetite separates concentrates (Table S1). Moreover, theassociation of the (Co, Mn, Ni) (OH)2 solid-solution phases with microtextures resembling fossilized
microorganisms coated by (Co–Mn–Ni) (OH)2 (Figure 10f, white arrows) and the occurrence of isolated
and aggregates of bacterio-morphic Fe–Ni–serpentine (Figure 10g) suggest the potential eﬀect of
microorganisms in the Vermio mineral-forming processes. Currently, much research interest has
focused on a consortium of microorganisms of varying morphological forms, producing enzymes,
which are considered to be a powerful factor to catalyze redox reactions, and act as nucleation sites
for the precipitation of secondary minerals [68–70]. In general, the presence of organic matter andmicroorganism traces in highly tectonized zones of the metamorphosed Fe–Ni-laterites may suggest
direct involvement of microorganisms, and have created appropriate conditions for metal bio-leaching
and bio-mineralization. In general karst-type Fe–Ni-laterites, where dominant Fe minerals are goethite± hematite [54,55,71], may display REE contents ranging from a few decades to thousands of more than
5000 ppm (Table S1) [54], depending on the site of the Ni-laterite samples along vertical proﬁles andthe post-deposition processes aﬀected their composition. The epigenetic mobilization of REE, alongwith Fe, Mn, Ni, Co, under reducing and acidic conditions and, subsequently, re-deposition underalkaline conditions at the lowest parts of contacts between laterites of Central Greece (Lokris) andcarbonate rocks has been well established as well [54–56,71]. Thus, the potential depletion of Ga andAl, and elevated Mn, Co, Zn, and Fe ± REE contents has been recorded in the above large Fe–Ni-laterite
deposits, but in small Fe–Ni-laterite occurrences in Northern Greece, strongly transformed during
post-magmatic metamorphism, the dominant Fe mineral is depleted magnetite, due probably to theirsmall size, facilitating potential changes during sub-solidus reactions. Therefore, the dominanceof magnetite in Fe–Ni-laterite ores is not an encouraging factor for the exploration of Ni sourcesof laterite-type.
5.3. Implications of the Trace Element Distribution for the Origin of Magnetite Mineralization
Despite the potential re-distribution of trace elements in magnetite separate-bearing deposits,disseminated magnetite separates from maﬁc–ultramaﬁc magmatic rocks of ophiolite complexes
exhibit much higher Ti and V contents compared to hydrothermal magnetite, and they are comparable
to magnetite separates from coastal sands derived from the weathering of calc-alkaline volcanic rocks(Table 5). The massive hydrothermal magnetite separates of skarn-type diﬀers from the disseminatedmagnetite separates from porphyry deposits, in terms of the lower Ba, Bi, Cr, Ni, Co, V, Sr, Sc, Rb,
ΣREE, Y, Hf, Th, U, and Zr contents (Table S1). Also, Ga is higher in magnetite separates associatedwith calc-alkaline rocks (38–59 ppm) compared to that related to ophiolites (2–21 ppm Ga) (Table
S1). The highest Ga contents were recorded in magnetite separates from the Skouries porphyry-Cu
deposit (40–59 ppm), skarn-type from Lavrion mine (38–39 ppm), and coastal sand from Kos Island (42ppm), all related with calc-alkaline rocks. A potential order of the average Ga content in magnetite
separates concentrates, from low to high, is from Ni-laterites, hydrothermal (ophiolites) associated withsulphides, ultramaﬁc lavas (ophiolites), hydrothermal (ophiolites) associated with apatite, magmatic(norite gabbros), skarn-type (calc-alkaline), coastal sands (calc-alkaline volcanic), and porphyry-Cu
deposits (hydrothermal calc-alkaline), resulting in a discrimination using a plot of (V + Ti) versus Ga
content (Figure 12).
Magnetite separates associated with Fe–Cu–Ni–Co sulphides with dominant minerals pyrrhotite,chalcopyrite, and pentlandite exhibit relatively high Cu and Zn, due to incomplete separation from theassociated sulphides, but it is REE-, Ti-, and V-depleted. Although an originally primary magmaticorigin for these sulphides is not precluded, characteristics of the highly transformed pure magnetite

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separates (Table 1 and Table S1) at Pindos may indicate that potential original magmatic features havebeen overprinted by a low-level hydrothermal circulation process.
The plot of the Eu/Eu* ratio versus Ga content shows a less negative anomaly Eu/Eu* (average 0.85)
for large Fe–Ni-laterite deposits [56,72,73], bauxite laterites and bauxites (63 and 0.69, respectively) [74],and the most negative Eu/Eu* anomaly (average 0.32), for the magnetite separates from Fe–Ni-laterites
of Northern Greece (Figure 12e; Table 5). Assuming that the presence of Eu anomaly, caused by
the ability of Eu to exist in either Eu2+ or Eu3+ states, the Eu/Eu* anomalies seem to be eﬀectedby physico/chemical conditions, such as a redox potential [74]. Diagenetic re-mobilization of Eu ispossible under conditions of reduction to Eu2+ at low oxidation potential, as is suggested by thepresence of organic matter [73–76]. In addition, the increasing order of the average Ga content from the
Fe–Ni-laterites to bauxite laterites and bauxites (Table 5; Figure 12e) is in agreement with the preferenceof Ga to high-Al ores. The average Ga content for karst bauxite deposits is 58 ppm with a range fromless than 10 to 812 ppm, because the close geochemical aﬃnity of Ga to Al enables Ga to substituteeasily in rock-forming alumino-silicates [77]. Such a preference Ga with high-Al chromitites, as welland the structural incorporation of Ga into chromite lattice, is evidenced by the progressive and linear
increase of Al or decreasing Cr/(Cr + Al) atomic ratio [78,79].
5.4. Implications of Magnetite Separates for Exploration of Precious Metals in Porphyry-Cu Deposits
Alkaline or K-rich calc-alkaline porphyry deposits worldwide represent signiﬁcant gold resources,
owing to their large sizes. They are derived from hydrous calc-alkaline magmas and are geneticallylinked to magmatic and hydrothermal processes. Despite their common origin in a supra-subduction
environment, platinum-group elements (PGEs) have been recorded only in certain porphyry-Cu
deposits. During the last decades, signiﬁcant Pd and Pt contents were described in the Cordillera of
British Columbia, in the Balkan–Karpathian system, including the Skouries deposit, in the Philippinesand elsewhere [4,12,13,80,81]. The estimated Pd, Pt, and Au potential for porphyry deposits, combinedwith the association of merenskyite [(Pd, Pt) (Te, Bi)2 )] with major Cu minerals (bornite and chalcopyrite),are considered to be encouraging economic factors for Pd and Pt, as by-products, the Au being themain precious metal product [13,80,81].
It has been suggested that ﬁrst-stage arc magmas tend to generate Cu-rich, relatively Au-poor
porphyry systems, while relatively Au-rich residue is left in the lower crust and lithospheric mantle
after ﬁrst-stage arc magmas [82]. Subsequently, a second stage of melting, due probably to lithosphericthickening, thermal rebound, mantle lithosphere delamination, or lithospheric extension, may generatemagmas with relatively high Au/Cu ratios [82–84] and PGE enrichments [84–87]. Critical requirements
for a signiﬁcant base/precious metal potential in porphyry Cu + Au + Pd ± Pt deposits are considered tobe the geotectonic environment, controlling the precious/base metal endowment in the parent magma(metasomatized asthenospheric mantle wedge) [86,87]. It is indicated by signiﬁcant Cr, Co, Ni, and Re
contents, the oxidized nature of parent magmas, that facilitate the capacity for transporting suﬃcient
Au and PGE, and the degree of evolution of the mineralized system, as is exempliﬁed by the enrichmentof (Pd + Pt) versus (Os + Ir + Ru) and LREE versus HREE [4,86,87]. Arc magmas with high potential toproduce hydrothermal systems with ideal chemistry for transporting precious metals (Cu + Au + Pd ±Pt deposits) are characterized by f O2 more than two log units above fayalite-magnetite-quartz (FMQ),
where f O2 is oxygen fugacity and FMQ is the fayalite–magnetite–quartz oxygen buﬀer [88]. Since theoxidized nature of parent magma is connected with the ability to produce a magmatic–hydrothermal
system with ideal chemistry that facilitates the capacity for transporting suﬃcient precious metals,
abundant magnetite (reaching up to 10 vol.% in the Skouries deposit) linked to pervasive potassic andpropylitic alteration type [13], in contrast to “reduced” porphyry Cu–Au deposits, lacking primary
hematite, magnetite separates, and sulphate minerals [89], is considered to be a characteristic feature of
the Pd-bearing porphyry-Cu deposits. Thus, the abundance of magnetite separates is a characteristicfeature of all porphyry-Cu–Au–Pd deposits, as indicated by unexpected high-grade Cu–(Pd + Pt)(up to 6 ppm) mineralization at the Elatsite porphyry deposit, which is found in a spatial association

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with the Chelopech epithermal deposit (Bulgaria) and the Skouries porphyry deposit, Greece [4,12].
Disseminated hydrothermal magnetite separate is often associated with ﬁne (less than 10 μm) crystalsof U-thorite and zircon within veinlets of quartz (Figure 3c), which is consistent with the elevated Th,
U, and Zr contents (Table S1).
The enrichment in Cr, Co ± Ni is a common feature, recorded in the Skouries, Elatsite, andother porphyry deposits of the Balkan Peninsula, characterized by signiﬁcant Pd and Pt contents, incontrast to the porphyry-Cu ± Mo deposits of Russia and Mongolia with lower (Pd + Pt) content(less than 25 ppm Cr [49,90]) supports the origin of their parent magma from an enriched mantle
source [91,92]. In addition, a salient feature is the evolved geochemical signature of the Skouries andElatsite porphyry-Cu deposits, as indicated by the negative trend between (Pd + Pt) and Cr contents(Figure 14). The application of stable isotope analyses of oxygen (δ18 O = 4.33–9.45%) and hydrogen(δD = −110% to −73%) on quartz veins from various drill holes and depths may provide evidence
to constrain the origin of ﬂuids trapped in quartz veins [85]. Although any systematic variation of
the isotopic data with depth is uncertain, it seems likely that the samples with higher Cr contents
correspond to lower δ18 O values, while those with lower Cr contents correspond to higher δ18 O
values. This isotope variation, coupled with the higher (Pd + Pt) content Pd/Pt ratios corresponding tohigher δ18 O values and lower Cr contents, may point toward the more evolved mineralizing system(Figure 14) [86].
6. Conclusions
Texture, mineral chemistry and geochemical data on massive and disseminated magnetite
separates from a wide range of geotectonic settings and rock types led us to the following conclusions:
•

•

•

•
•

•

Magnetite separates of hydrothermal origin associated with calc-alkaline rocks can be distinguishedfrom that associated with ophiolites in terms of their higher Ga content, while magnetite separates
of magmatic origin are characterized by the highest (V + Ti) content.
Hydrothermal magnetite separates from skarn-type (Lavrion mine) diﬀers compared to thedisseminated separates from porphyry-Cu deposit (Skouries) in terms of the higher Ba, Bi, Cr, Ni,
Co, V, Sr, Sc, Rb, ΣREE, Y, Hf, Th, U, and Zr contents in the latter.
Although an initial primary magmatic origin for the Fe–Cu–Ni–Co sulphides in the Pindos
ophiolite is not precluded, geochemical characteristics (REE-, Ti-, and V-depleted) of the highly
transformed magnetite separates may indicate that potential original magmatic features havebeen overprinted by low-level hydrothermal circulation processes.
Similarity between the REE patterns, including the clear Eu negative anomaly of magnetite
separates from coastal black sand of Kos Island, may provide evidence of an andesitic source.
A major requirement controlling the Pd and Pt potential of porphyry-Cu deposits is probably theoxidized nature of parent magmas that facilitate the capacity for transporting suﬃcient Pd and Pt,
and the crystallization of abundant magnetite.
Abundance of magnetite in porphyry-type deposits may be used to discriminate between
porphyry-Cu–Au–Pd–Pt and porphyry-Cu–Mo deposits lacking precious metals, providing
evidence for the exploration of precious metals in porphyry systems.

Supplementary Materials: The following are available online at http://www.mdpi.com/2075-163X/9/12/759/s1,
Table S1: Major and trace element contents in representative magnetite separates from various geotectonic settings(Greece).
Author Contributions: D.G.E. collected the samples, provided the ﬁeld information, and performed all types ofanalyses. Both authors (D.G.E. and M.E.-E.) contributed to the elaboration and interpretation of the data. The
writing of the paper was completed by M.E.-E.
Funding: This research was funded by the National and Kapodistrian University of Athens (NKUA) (Grant No.
KE_11730).

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Minerals 2019, 9, 759

Acknowledgments: Many thanks are expressed to the two anonymous reviewers for their constructive criticism
and suggestions on an earlier draft of the manuscript. Vassilis Skounakis, University of Athens, is thanked for his
assistance with the SEM/EDS analyses.
Conﬂicts of Interest: The authors declare no conﬂicts of interest.

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minerals
ArticleMineralogical and Geochemical Constraints on theOrigin of Maﬁc–Ultramaﬁc-Hosted Sulphides:
The Pindos Ophiolite Complex
Demetrios G. Eliopoulos 1 , Maria Economou-Eliopoulos 2, *, George Economou 1 andVassilis Skounakis 2
1
2*

Institute of Geology and Mineral Exploration (IGME), Sp. Loui 1, Olympic Village,
GR-13677 Acharnai, Greece; eliopoulos@igme.gr (D.G.E.); Georgeoik7@gmail.com (G.E.)
Department of Geology and Geoenvironment, University of Athens, 15784 Athens, Greece;
vskoun@geol.uoa.gr
Correspondence: econom@geol.uoa.gr

Received: 16 April 2020; Accepted: 13 May 2020; Published: 18 May 2020

Abstract: Sulphide ores hosted in deeper parts of ophiolite complexes may be related to either primarymagmatic processes or links to hydrothermal alteration and metal remobilization into hydrothermal
systems. The Pindos ophiolite complex was selected for the present study because it hosts both
Cyprus-type sulphides (Kondro Hill) and Fe–Cu–Co–Zn sulphides associated with magnetite(Perivoli-Tsoumes) within gabbro, close to its tectonic contact with serpentinized harzburgite,
and thus oﬀers the opportunity to delineate constraints controlling their origin. Massive Cyprus-typesulphides characterized by relatively high Zn, Se, Au, Mo, Hg, and Sb content are composed of pyrite,
chalcopyrite, bornite, and in lesser amounts covellite, siegenite, sphalerite, selenide-clausthalite,
telluride-melonite, and occasionally tennantite–tetrahedrite. Massive Fe–Cu–Co–Zn-type sulphides
associated with magnetite occur in a matrix of calcite and an unknown (Fe,Mg) silicate, resemblingMg–hisingerite within a deformed/metamorphosed ophiolite zone. The texture and mineralogicalcharacteristics of this sulphide-magnetite ore suggest formation during a multistage evolutionof the ophiolite complex. Sulphides (pyrrhotite, chalcopyrite, bornite, and sphalerite) associatedwith magnetite, at deeper parts of the Pindos (Tsoumes), exhibit relatively high Cu/(Cu + Ni) andPt/(Pt + Pd), and low Ni/Co ratios, suggesting either no magmatic origin or a complete transformationof a preexisting magmatic assemblages. Diﬀerences recorded in the geochemical characteristics,
such as higher Zn, Se, Mo, Au, Ag, Hg, and Sb and lower Ni contents in the Pindos compared to the
Othrys sulphides, may reﬂect inheritance of a primary magmatic signature.
Keywords: sulphides; ophiolites; ultramaﬁc; selenium; gold; Pindos

1. Introduction
Traditionally, the sulphide mineralization associated with ophiolite complexes is that ofCyprus-type volcanogenic massive sulphide (VMS) deposits. They may be derived from the interactionof evolved seawater with maﬁc country rocks, under greenschist facies metamorphic conditions andsubsequent precipitation on and near the seaﬂoor, when ore-forming ﬂuids are mixed with cold
seawater [1–5]. They are associated with basaltic volcanic rocks and are important sources of baseand trace metals (Co, Sn, Se, Mn, Cd, In, Bi, Te, Ga, and Ge) [1]. Massive sulphide deposits have been
described in the Main Uralian Fault Zone (Ivanovka and Ishkinino deposits), southern Urals; they are
maﬁc–ultramaﬁc-hosted VMS deposits and show mineralogical, compositional, and textural analogies
with present-day counterparts on ultramaﬁc-rich substrates [6]. Recently, an unusual associationof magnetite with sulphides of Cyprus-type VMS deposit was described in Ortaklar, hosted in the
Minerals 2020, 10, 454; doi:10.3390/min10050454

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Koçali Complex, Turkey, which is part of the Tethyan Metallogenetic Belt [7]. Additionally, the largest
magnetite deposit in a series of apatite and sulphide-free magnetite orebodies hosted in serpentinites of
Cogne ophiolites, in the Western Alps, Italy, is characterized by typical hydrothermal compositions [8].
Although the Fe–Cu–Ni–Co-sulphide mineralization was initially considered an unusual type inophiolite complexes, several occurrences have been located, like those in pyroxenite cumulates of theOregon ophiolite [9], in dunites associated with chromitites of the Acoje ophiolite, Philippines [10],in layered gabbros of the Oman ophiolite [11], in dunites of the upper mantle–crust transition zoneof the Bulqiza (Ceruja, Krasta), Albania ophiolite [12], Shetland (Unst), UK ophiolite [13], and the
Moa-Baracoa ophiolitic massif (Cuba) [14]. On the basis of magmatic texture features and steep positivechondrite-normalized Platinum-Group Elements (PGE) patterns, sulphide mineralization of that typehas been interpreted as reﬂecting the immiscible segregation of sulphide melts [9,13,14]. Moreover, theoccurrence of Fe–Ni–Cu±Zn-sulphide mineralization (with dominant minerals pyrrhotite, chalcopyrite,and minor pentlandite) in mantle serpentinized peridotites and maﬁc to ultramaﬁc rocks of ophiolitecomplexes of Limassol, Cyprus, Othrys (Eretria) in Greece, Pindos (Tsoumes) ophiolite and elsewhere
has been the topic of research for extensive studies [15–20]. On the other hand, texture and geochemical
characteristics, including PGE contents, and a very low partition coeﬃcient for Ni and Fe between
olivine and sulphides are inconsistent with sulphides having an equilibrium with Ni-rich host rocks
at magmatic temperature [16,18]. Although the initial magmatic origin is not precluded, presentcharacteristics of the highly transformed ore at the Eretria (Othrys) area may indicate that the magmaticfeatures have been lost or that metals were released from the host rocks by a low-level hydrothermal
circulation process [16]. Fe–Cu–Zn–Co–Ni mineralization is also reported in seaﬂoor VMS deposits
from modern oceans (as well as in their possible analogues on several ophiolites on land, e.g., Urals)
indicating that these deposits can be formed by purely hydrothermal processes [19].
Despite the extensive literature data on a diverse array of sulphide mineralizations, sulphide ores
hosted in maﬁc–ultramaﬁc ophiolitic rocks are characterized by structure as well as mineralogical andgeochemical features, suggesting either magmatic origin or links to serpentinization processes andmetal remobilization from primary minerals into hydrothermal systems. The present study is focusedon some new SEM/EDS and geochemical data on Cyprus-type and Fe–Cu–Co–Zn-type sulphides
hosted in deeper parts of the Pindos ophiolite complex, aiming to improve our understanding of thefactors controlling trace element incorporation into sulphide minerals and their origin.
2. Materials and Methods
2.1. Mineral Chemistry
Polished sections (20 samples) from sulphide ores were carbon-coated and examined by a scanningelectron microscope (SEM) using energy-dispersive spectroscopy (EDS). The SEM images and EDS
analyses were carried out at the University of Athens (NKUA, Athens, Greece), using a JEOL JSM
5600 scanning electron microscope (Tokyo, Japan), equipped with the ISIS 300 OXFORD automatedenergy-dispersive analysis system (Oxford, UK) under the following operating conditions: accelerating
voltage 20 kV, beam current 0.5 nA, time of measurement (dead time) 50 s, and beam diameter 1–2 μm.
The following X-ray lines were used: FeKα, NiKα, CoKα, CuKα, CrKα, AlKα, TiKα, CaKα, SiKα,
MnKα, and MgKα. Standards used were pure metals for the elements Cr, Mn, Mo, Ni, Co, Zn, V,
and Ti, as well as Si and MgO for Mg and Al2 O3 for Al.
2.2. Whole Rock Analysis
The studied sulphide samples were massive and disseminated mineralizations, derived fromlarge (weighing approximately 2 kg) samples, which is necessary to obtain statistical representative
trace element distribution in sulphide ores. They were crushed and pulverized in an agate mortar.
Major and minor/trace elements were determined at the SGS Global—Minerals Division Geochemistry
Services Analytical Laboratories Ltd., Vancouver, BC, Canada. The samples were dissolved using

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sodium peroxide fusion, combined Inductively Coupled Plasma and Atomic Emission Spectrometry,
ICP-AES and Mass Spectrometry, ICP-MS (Package GE_ICP91A50). On the basis of the quality control
report provided by Analytical Labs, the results of the reference material analysis in comparison to
expected values, and the results from the multistage analysis of certain samples, showed an accuracy
and a precision of the method in good agreement with the international standard (<10%).
3. A Brief Outline of Characteristics for the Studied Sulphides
The Pindos ophiolite complex, of Middle to Upper Jurassic age, is located in the northern-westernpart of Greece (49◦ N, 21◦ E), lies tectonically over Eocene ﬂysch of the Pindos zone, and contains aspectrum of lavas from Mid-Ocean Ridge, MOR basalts through island arc tholeiites (IATs) to boninite
series volcanics (BSVs) [21–23]. Two tectonically distinct ophiolitic units can be distinguished: (a) the
upper unit (Dramala Complex), including mantle harzburgites, and (b) a lower unit, including volcanicand subvolcanic sequences at the Aspropotamos Complex (Figure 1). This complex consists of astructurally dismembered sequence of ultramaﬁc and maﬁc cumulate ophiolitic rocks, including
gabbros, which is locally underlain by sheets of serpentinite [21,22]. Sulphide mineralization in the
Pindos ophiolite complex is located near the Aspropotamos dismembered ophiolite unit, belonging to
the lower ophiolitic unit of the complex and includes the following.

Figure 1. Simpliﬁed geological map of the southern part of the Pindos ophiolite, showing the Perivoli(Tsoumes–Kondro) area (modiﬁed after [22,24]).

A volcanic and subvolcanic sequence composed mainly of basalts and basaltic andesite pillow
lavas ranging from high to low Ti aﬃnity [20–24]. The diﬀerent magmatic groups may have been
derived from diﬀerent mantle sources and/or various degrees of partial melting [20–24]. Massive
Cyprus-type sulphide ore occurrences in the form of small lenses (maximum 4 × 40 m), are located in an
abandoned mine, at the Kondro Hill, very close to the village of Perivoli (Figure 1). The estimated ore
potential is about 10,000 tons with an average 6.6 wt.% Cu and 9.4 wt.% Zn [25]. They occur on the top
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of diabase (massive or pillow lavas) and are directly overlain by metalliferous (Fe-Mn-oxide-bearing
sediments). Due to the tectonic disruption of the Aspropotamos unit, the spatial association between
massive and stockwork disseminated mineralization is unclear.
Several sub-vertical veins of quartz with veinlets and a brecciated pipe-shaped diabase dike(stockwork ore zone) have been described in the Neropriona area of the Aspropotamos unit, Kondro
Hill, with disseminations of pyrite + chalcopyrite, and dominant mineral-altered plagioclase andclinopyroxene, penninite, kaolinite, quartz, epidote, and calcite [20,25–27]. Small irregular to lens-likeoccurrences (4 × 1.5 m) of massive Fe–Cu–Zn–Co-type sulphide mineralization associated withmagnetite are exposed at the Perivoli (Tsoumes) Hill (Figure 1). These are hosted within gabbro,
close to its contact with serpentinized harzburgite [20,25], consisting of pyrrhotite, pyrite, chalcopyrite,and sphalerite associated with magnetite. The contact between ore and hosting rock is not sharp,
appearing as irregular nets of veinlets. Rounded fragments of highly altered rock and massive
fragments of sulphide ore are broadly parallel to the shear plane of a thrust fault.
4. Mineralogical Features
4.1. Cyprus-Type Sulphides
The massive ore is mainly composed of pyrite, chalcopyrite, bornite, and in lesser amounts
covellite, siegenite, sphalerite, and clausthalite, while pyrrhotite is lacking (Figure 2). Chalcopyrite,bornite, and sphalerite occur in at least two diﬀerent generations. Pyrite grains vary from euhedral tosubhedral and rarely framboidal. Textural relationships indicate that early pyrite, commonly occurring
as large crystals but often exhibiting dissolution, is extensively penetrated and replaced by ﬁne-grained
chalcopyrite, bornite, and sphalerite in a matrix of quartz (Figure 2). Copper-bearing sphalerite,with up to 3.6 wt.% Fe, 4.2 wt.% Cu, and 1.7 wt.% Bi, occurs within pyrite crystals and/or cements minor
chalcopyrite and pyrite (Figure 2b,e,f; Table 1). Pyrite is extensively replaced by intergrowths between
chalcopyrite or bornite and Fe-poor sphalerite (Figure 2b,d; Figure 3) and occasionally contains Co(Table 1). Fine-grained intergrowths of framboidal or colloform pyrite-bornite, occurs in a matrixof quartz (Figure 2g,h). Fine-grained chalcopyrite or bornite are often found in cross-cutting veins,
hosting selenides (mainly clausthalite, PbSe) (Table 1), the telluride mineral melonite (NiTe2 ), gold,
galena, and barite [26]. Furthermore, present investigation reveals the formation of aggregates ofsecondary minerals, occurring as characteristic crusts on bornite surfaces (Figure 2i–l). These minerals
are present-day grown minerals, on the surface of polished sections of sulphide ore, exposed to air,under room conditions (20–25 ◦ C) and moderate air humidity (atmospheric water). Gold, as inclusions
in chalcopyrite reaching a maximum size of 20 μm with Ag contents of up to 9 wt.%, is a rare componentof the ores [26]. Additionally, we observed the presence of submicroscopic gold, i.e., <1 μm and thus
invisible under an optical microscope, in grains of pyrite, chalcopyrite, and bornite, that increases
with decreasing crystal size, reaching contents up to 7.7 ppm Au in pyrite, 8.8 ppm Au in very ﬁneintergrowths between pyrite and sphalerite, and 17.3 ppm Au in ﬁne intergrowths between pyrite andbornite [20].

222Minerals 2020, 10, 454

Figure 2. Backscattered electron (BSE) images representative of various morphological forms andtextures of pyrite associated with chalcopyrite and sphalerite in a matrix of quartz (a–d); euhedral pyrite
displaying erosion and replacement by chalcopyrite–sphalerite intergrowths (b); a close intergrowth
between pyrite and bornite (d,e); selenides (clausthalite) as inclusions within bornite (f); ﬁne-grainedintergrowths of framboidal or colloform pyrite-bornite, in a matrix of quartz (g,h); replacement ofbornite by neo-formed intergrowths of Cu minerals (i–l), which are Cu-enriched and Fe-depleted(Table 1). Scale bar: 200 μm (a); 50 μm (b,f,g,i); 20 μm (e); 10 μm (c,d,h,j,k,l). Abbreviations: py = pyrite;
cp = chalcopyrite; sph = sphalerite; bn = bornite; PbSe = clausthalite; qz = quartz; f.py = framboidal
forms of pyrite with tiny inclusions of Cu minerals; ge-sp = intergrowths of neo-formed Cu minerals
with composition corresponding to geerite and spionkopite.

223

3
42.8
2.1
n.d.
n.d.
n.d.
n.d.
53.9
99.3

14
3.2
70.6
4.6
n.d.
n.d.
n.d.
n.d.
n.d.
21.6
100

224

Pyrrhotite
26
58.9
n.d.
0.8
1.4
39.8
100.6

25
61.0
n.d.
n.d.
n.d.
39.4
100.4

Fe
Cu
Zn
Bi
S
Total

16
0.8
82.4
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
26.2
99.4

17
3.2
73.2
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
24.2
100.6

18
1.4
73.2
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
25
99.6

19
0.6
76.2
0.9
n.d.
n.d.
n.d.
n.d.
n.d.
23
100.7

27
59.7
n.d.
n.d.
n.d.
40
99.7

Pyrite
28
29
44.9
44.5
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
55.3
54.2
100.2
99.7
30
32.0
31.4
2.3
n.d.
34.3
100

20
3.2
4.6
n.d.
21.6
70.5
n.d.
n.d.
n.d.
n.d.
99.9

Chalcopyrite
31
32
30.6
30.4
33.8
33.6
n.d.
n.d.
n.d.
n.d.
35.9
10
100.3
99.6

Sulphides associated with Magnetite

15
2.2
71.1
1.8
n.d.
n.d.
n.d.
n.d.
n.d.
23.9
99

22
2.1
2.3
n.d.
24.8
69.8
n.d.
n.d.
n.d.
n.d.
99

Sphalerite
33
34
7.5
6.4
n.d.
n.d.
57.8
58.6
1.4
1.9
33.1
33.2
99.8
100.1

Tsoumes

21
2.2
1.9
n.d.
25.1
71
n.d.
n.d.
n.d.
n.d.
100.2

12
11.9
62.2
n.d.
n.d.
n.d.
0.6
25.3
100

23
0.9
1.5
n.d.
1.4
n.d.
1.9
13.8
80.4
n.d.
99.9

24
1.2
0.9
n.d.
0.5
n.d.
3.8
12.4
80.9
n.d.
99.7

Tellurides-Melonite

KondroBornite
10
11
11
11.7
60.6
60
n.d.
n.d.
n.d.
n.d.
1.2
n.d.
n.d.
n.d.
27.1
27.8
99.9
99.5

Selenides-claousthalite

Sulphides of Cyprus Type
Sphalerite
Chalcopyrite
4
5
6
7
8
9
3.6
1.9
1.8
28.7
29.9
29.2
4.2
1.2
0.4
33.6
34.0
35.3
57.7
62.7
63.3
1.6
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
1.7
1.4
1.5
0.8
1.2
n.d.
n.d.
n.d.
n.d.
n.d.
n.d.
1.1
32.8
33.3
32.9
35.3
35.4
32.9
100
100.5
99.9
100
100.4
99.8

Epigenetic Cu-minerals (Figures 2 and 3)

13
3.8
70.8
n.d.
n.d.
n.d.
0.3
n.d.
n.d.
24.5
99.4

1
44.6
n.d.
n.d.
n.d.
1.2
n.d.
54.1
99.9

Pyrite
2
43.7
0.5
n.d.
1.2
n.d.
n.d.
53.9
99.3

wt%
Fe
Cu
Zn
Se
Pb
Co
Ni
Te
S
Total

MineralMineral
wt%
Fe
Cu
Zn
Co
Bi
Se
S
Total

1 = Fe3.2Bi0.02Cu0.05S6.8
2 = Fe3.1Co0.1Cu0.03S6.7
3 = Fe3.0Cu0.14S6.8
4 = Zn4,3Fe0.3Bi0.04Cu0.3S5.0
5 = Zn4,7Fe0.2Bi0.03Cu0.09S5.0
6 = Zn4.8Fe0.2Bi0.04S5.0
7 = Cu2.4Fe2.4Zn0.02S5.1
8 = Cu2.5Fe2.4Bi0.02S5.1
9 = Cu2.5Fe2.4Bi0.04S5.1
10 = Cu4.8Fe1.0Bi0.03S4.3
11 = Cu4.8Fe1.0S4.2
12 = Cu4.9Fe1.1Se0.04S4.0
13 = Cu5.7Fe0.4S4.1
14 = Cu5.8Fe0.3Zn0.4S3.5
15 = Cu5.8Fe0.2Zn0.2S3.9
16 = Cu6.1Fe0.0S3.8
17 = Cu5.9Fe0.3Co0.3S3.9
18 = Cu5.9Fe0.12S4.0
19 = Cu6.2Fe0.06Zn0.07S3.7
20 = Pb4.6Se3.7Fe0.8Cu1.0
21 = Pb4.7Se4.4Fe0.5Cu0.4
22 = Pb4.6Se4.3Fe0.5Cu0.5
23 = Te6.97Ni2.6Fe0.2Cu0.2Se0.2
24 = Te7,17Ni2.4Fe0.3Cu0.6Se0.07
25 = Fe4.7S5.1
26 = Fe4.6Zn0.05Bi3.03S5.3
27 = Fe4,6S5.4
28 = Fe3.2S6.8
29 = Fe3.2S6.8
30 = Fe2.6Cu2.3Zn1.6S4.9
31 = Fe2.5Cu2.4S5.1
32 = Fe2.5Cu2.4S5.1
33 = Zn4.3Fe0.6Bi0.03S5.0
34 = Zn4.4Fe0.6Bi0.04S5.0

Table 1. Representative SEM/EDS analyses of minerals from the Pindos sulphide ores. (n.d.: Lower than detection limit).

Minerals 2020, 10, 454

Minerals 2020, 10, 454

Figure 3. Backscattered electron (BSE)images showing bornite, partially replaced by neo-formedintergrowths of epigenetic minerals resembling geerite or spionkopite, in a matrix of quartz (a) and thesingle scanning for Fe (b), Si (c), and S (d). White arrows indicate the Fe depletion in the present-day
formed intergrowths of high Cu minerals. Symbols, as in Figure 2. Scale bar for b, c, d as in 3a (20 μm)

4.2. Breccia Pipe
Disseminated pyrite, minor chalcopyrite, and sphalerite occur mostly in vesicles ﬁlled by quartz,
kaolinite, chlorite, and epidote, within brecciated pipe-form diabase (a discharge pathway) underlying
the Kondro massive ore [27]. Samples (n = 10) of pyrite separates from the diabase breccia have showna limited range for δ34 S values from +1.0 to +1.5% [27].
4.3. Massive Fe–Cu–Co–Zn-Type Sulphides Associated with Magnetite
The sulphide ore is mainly composed of pyrrhotite, while pyrite, chalcopyrite, sphalerite, malachite,
and azurite are present in lesser amounts. Pure magnetite, often forming a network texture, is associatedwith sulphides, either as massive ore with inclusions of sulphides (chalcopyrite, pyrite, and pyrrhotite),or as individual grains dispersed within sulphide ore (Figure 4). A characteristic feature of magnetite
is its elongated and curved form and its textural relationship with the sulphides. Pyrrhotite, whichis the most abundant sulphide, is followed by chalcopyrite and sphalerite, all showing an irregularcontact with the magnetite, that often occurs surrounding sulphides (Figure 4a–c). Additionally,a salient feature is the occurrence of an (Fe/Mg) phyllosilicate associated with sphalerite, chalcopyrite,and magnetite (Figure 4d–f; Figure 5).

225

Minerals 2020, 10, 454

Figure 4. Backscattered electron (BSE) images showing intergrowths between pyrrhotite, sphalerite,and magnetite in a matrix of calcite (a); elongated curved crystals of magnetite and inclusions within
pyrrhotite (b); intergrowths between pyrrhotite, sphalerite, chalcopyrite, and magnetite in a matrix ofcalcite (c); magnetite associated with sphalerite in a matrix of calcite and (Fe,Mg) silicate (d); sphalerite
adjacent to chalcopyrite with inclusions of (Fe,Mg) silicate (e); transitional contact between chalcopyriteand sphalerite and their intergrowths with magnetite (f) Scale bar: 100 μm (a,c,d); 50 μm (b); 20 μm (e);10 μm (f). Symbols: cal = calcite; Fe-Mg-Si = (Fe,Mg) silicate, and as in Figure 2.

226

Minerals 2020, 10, 454

Figure 5. Backscattered electron (BSE) images showing intergrowths of pyrrhotite, pyrite, chalcopyrite,and sphalerite in a matrix of calcite and an unknown (Fe–Mg) silicate (a), Table 2; single scanning for S(b), Fe (c), Cu(d), and Zn (e). Scale bar for b, c, d as in 5a (50 μm). Symbols, as in Figures 2 and 4.
Table 2. Representative SEM/EDS analyses of (Fe,Mg) silicates from the Tsoumes massive sulphides.
wt %Tsoumes(Fe,Mg)-Silicate

Laramie Complex [28]Hisingerite

SiO2
Al2 O3Cr2 O3Fe2 O3t
MnO
MgO
NiO
CaO

35
0.9
n.d.
42.3
n.d.
11.9
n.d.
n.d.

33.2
1.1
n.d.
43.2
n.d.
12.1
n.d.
n.d.

42.7
1.2
0.03
35.83
0.15
7.46
0.05
0.81

37.8
0.03
0.04
46.94
0.49
2.15
0.01
0.39

Total

90.1

89.6

88.22

87.85

227

Minerals 2020, 10, 454

5. Geochemical Characteristics of Sulphides
Massive sulphide ores from the Kondro Hill exhibit uncommon high contents in Au (up to
3.6 ppm), Ag (up to 56 ppm), Se (up to 1900 ppm), Co (up to 2200 ppm), Mo (up to 370 ppm),
Hg (up to 280 ppm), Sb (up to 10 ppm), and As up to 150 ppm, which are much higher than thosein the Fe–Cu–Zn–Co-type sulphide hosted in ultramaﬁc parts of the complex, along a shearing zone,
close to a contact with gabbros (Table 3), as well as within brecciated pipe-form diabase (a discharge
pathway) underlying the Kondro massive ore [20]. Major and minor elements, such as Fe, Cu, and Zn,
are hosted in sulphides (pyrrhotite and pyrite, chalcopyrite, bornite, and epigenetic high Cu minerals,
and sphalerite, respectively). Magnetite and Fe silicates (Tables 1 and 2), selenides, tellurides, gold,
galena, and barite are occasionally present, but Mo-bearing minerals were not identiﬁed.
Major and trace elements in massive sulphide ores of Cyprus- and Fe–Cu–Co–Zn-type sulphides
from the Pindos ophiolite complex, along with those from comparable ophiolites such as the Othrys
and Troodos ophiolite complexes [29–33], are plotted in Figure 6. Although there are overlappingﬁelds, it seems likely that the Pindos sulphides can be distinguished by their higher Zn (Figure 6a,c)
and Co (Figure 6b) contents (Figure 6). The highest Se contents were recorded in the Pindos and the
Apliki ores (Cyprus) accompanied by Cu and Au contents (Figure 6d,e). A positive trend is also clear
between Au and As (Figure 6f).

228

Minerals 2020, 10, 454

Figure 6. Plots of major and trace elements of Cyprus-type (labeled as Cy) sulphide ores and
Fe-Cu ± Co ± Zn sulphides hosted in ultramaﬁc rocks (labeled as ultram.) from the Pindos (Pi), Othrys,
and Troodos ophiolite complexes. Although there are overlapping ﬁelds, the ores from the Pindos canbe distinguished for the presence of ores with higher Zn (a,c) and Co (b) contents in Cyprus-type ores.
The highest Se contents are recorded in the Pindos and the Apliki ores (Cyprus) accompanied by Cu andAu contents (d,e). A positive trend is also clear between Au and As (f). Data: Table 3; [3,7,16,18,20,29,30].

229

230

CyprusOthrysLimassol

Eretria

Tsoumes

Disseminated Diabase breccia

Aspropotamos

Kondro Massive ore

PindosDescription

Location

0.006
0.007

1.8
2.5
1.4
1.7

3.3
0.44
0.89

1.8
1.74
3.8
0.91
0.35
0.31

49.8
33.4
31.2
30.3

39.4
35.9
37.4

38.8
43.5
43.8
38.7
54.4
13.2

6.9
6.1
16.4
10.1
25.4
7.2
18.9
11.6
8.1

Cu

wt %

7.1
6.9

26.9
21.2
26.5
22.5
11.9
20.8
9.5
26.4
23.2

Fe

wt %

0.12
0.03
0.34
0.007
0.009
0.006

0.02
0.12
0.01

0.4
0.12
0.11
0.1

0.007
0.009

2.9
2.6
3.2
9.8
1.1
2.2
1
8.3
2.5

Zn

wt %

1500
1800
2300
3600
1700
500

2300
400
2000

600
1100
540
520

27
15

1400
1250
600
500
2200
910
870
280
1000

Co

40
610
70
40
10
12

40
30
46

80
150
70
58

<5
<5

50
80
310
40
80
91
76
36
370

Mo

35
5
110
45
44
6

6
7
8

130
350
95
41

<5
<5

170
400
1500
1900
850
350
1100
1900
1400

Se

47
12,800190
54,800
5600
730

2
<5
<5

20
27
20
11

25
25

140
150
90
64
40
140
10
64
200

As

1.1
1.3
3.3
3.6
3.2
3
2.1
1.9
1.9

Au

3.5
8.2
4.5
6.2
2.5
0.53

0.01
0.02
0.01

0.15
0.14
0.11
0.03

0.018
0.025

Symbol __: Pt contents < 10 ppb.

12,00017,1001400
48,50012,500
3500

2400
3700
2500

9
12
8
7

<5
<5

20
32
12
11
27
110
60
8
13

Ni

2
<1
<1
2
1.5
1

1
0.3
0.6

<1
<1
<1
2.1
8.2
4.5
0.91
0.35
0.31

10
30
11
16

<1
<1

40
70
170
280
60
48
128
280
168

Hg

6
7
8
5

<1
<1

56
34
38
32
39
35
40
38
48

Ag

ppm

18
73
10
94
13
6.6

0.2
0.2
0.2

1
0.8
0.9
0.4

0.3
0.6

9.5
7.1
8.1
10
2.1
6.6
1.3
10
1.4

Sb

0.6
0.5
0.96
0.66
0.21
0.53

8
9.5
0.61
13.5
7.35
7

1.04
12.33
1.25

0.015
0.01
0.015
0.013

>0,99
>0,99
>0,99
>0,99
0.93
0.49
0.78

__
__

0.014
0.026
0.02
0.022
0.012
0.12
0.07
0.028
0.013

Ni/Co

__
__

>0,99
>0,99
>0,99
>0,99
>0,99
>0,99
>0,99
>0,99
>0,99

Cu/(Cu +Ni)

Table 3. Geochemical characteristics of sulphide ores hosted in the Pindos, Othrys, and Troodos ophiolites. Data: present study [20].

__
__
__
0.95
0.5
0.61

0.99
0.61
0.86

0.12
0.87
0.22
0.87

0.95
0.97

__
__
__
__
__
__
__
__
__

Pt/(Pt + Pd)

Minerals 2020, 10, 454

Minerals 2020, 10, 454

6. Discussion
6.1. A Comparison between Magmatic Sulphides and Those Hosted in Maﬁc–Ultramaﬁc Ophiolites
It has been well established that the formation of magmatic deposits is related to the segregation ofsulphide melts by immiscibility from basaltic magmas, which are able to collect precious metals (PGE, Ag,Au) as well as other chalcophile elements (Se, Te, Bi, Pb, As, Sb) because of their high partition coeﬃcient(Dsulphide melt/silicate melt ) [34,35]. The compositions of the magmatic ores may be mainly controlled by(a) the metal abundances of the mantle source and the degree of partial melting [36,37], (b) the degreeof fractional crystallization and potential crustal contamination during magma ascent [38], and (c)interactions with magmatic-hydrothermal or metamorphic-hydrothermal ﬂuids [39–41]. Compositional
data obtained from a large number of magmatic sulphide deposits [42] show clear diﬀerences from the
Fe–Cu ± Co ± Zn ± Ni-sulphide mineralization hosted in ophiolite complexes.
The unusual Fe–Cu–Ni–Co type of sulphides hosted in maﬁc–ultramaﬁc ophiolitic rocks ischaracterized by varying structural, mineralogical, and geochemical features, which are not ofmagmatic origin. Sulphides hosted in the magmatic sequence of ophiolite complexes [9–14] havebeen interpreted as reﬂecting the immiscible segregation of sulphide melts, because they exhibit
magmatic texture features and steep positive chondrite-normalized PGE patterns [9–14]. However,the Pindos (Tsoumes) Fe–Cu–Co–Zn-sulphides, consisting mainly of pyrrhotite, chalcopyrite, bornite,and sphalerite and hosted within gabbro, close to its tectonic contact with serpentinized harzburgite,
diﬀer compared to the Othrys (Eretria) sulphides, which are located at the peripheries of podiform
chromite bodies hosted in serpentinized harzburgite and consist of pyrrhotite, chalcopyrite, and minor
Co-pentlandite. Furthermore, the Pindos massive sulphides diﬀer from those in the serpentinizedrocks of the Limassol Forest (Troodos), which are composed dominantly of troilite, maucherite,
pentlandite, chalcopyrite, bornite, vallerite, magnetite, minor sphalerite, graphite, molybdenite,
and gold. In addition, the sulphides in the Limassol Forest contain much higher Au and As, up to
8 and 62 ppm, respectively [15,18], compared to the Pindos sulphides (Table 3). Despite the above
diﬀerences between the Pindos, Othrys, and Limassol sulphides, they are all characterized by higher
Cu/(Cu + Ni) ratios (>0.99, 0.5–0.93, and 0.2–0.96, respectively) compared to most magmatic deposits,
having Cu/(Cu + Ni) ratios ranging from 0.05 to 0.14 [42]. In addition, the Ni/Co ratio in magmatic
deposits typically ranges from 15 to 50 [42], whereas the range of Ni/Co ratio is 1.0–12 for the Eretria,
0.6–14 for the Limassol, and 0.01–0.02 for the Pindos (Tsoumes) samples (Table 3).
Model calculations have shown that the relatively high Co tenor and low Ni/Co cannot be
explained by an earlier phase of fractional crystallization or sulphide segregation [43]. Thus, the higher
Ni and Co contents in sulphide occurrences associated with chromitite bodies in the Othrys (Eretria)
peridotites may suggest re-mobilization of Fe, Co, and Ni during hydrothermal alteration of peridotite
in the presence of aqueous H2 S, and precipitation of Ni–Co–Fe sulphides [44,45]. The presence ofgraphite-like material in chromitites associated with sulphides from the Othrys complex, along shear
zones that served as ﬂuid pathways through the chromitites [46,47], may support the mobilization andre-precipitation of Fe–Ni–Co sulphides. Furthermore, the occurrence of phosphides such as Ni–V–Co
phosphide [48], Mo3 Ni2 P1 + x [49], NiVP [50], and the associated sulphide V7 S8 [51] in chromitite
concentrates from the Othrys ophiolite are consistent with extremely low fO2 (reducing environment)
during serpentinization [51] and re-precipitation of sulphides.
Intergrowths between sulphides and magnetite often forming curved crystals, which reﬂect asimultaneous deposition, coupled with the occurrence of calcite and an unknown (Fe,Mg) silicate(Figure 4d–f) resembling Mg–hisingerite [52] may provide evidence for the conditions of the Pindos
sulphide deposition. Such a (Fe,Mg) silicate is unusual compared to the common presence of
Mg-enriched serpentine in other ophiolite complexes. It seems to be comparable with the secondary
phyllosilicates described in altered ferroan metaperidotite (“Oxide Body”), from the Laramie Complex(Laramie city, WY, USA), and serpentinites containing the Si-free minerals, such as brucite and
NiFe alloy (awaruite). Such secondary phyllosilicates contain approximately equal amounts of231

Minerals 2020, 10, 454

end-members of the serpentine [(Mg,Fe2+ )3 Si2 O5 (OH)4 ] and hisingerite [Fe3+ 2 Si2 O5 (OH)4 .nH2 O] [28].
The substitution of Fe3+ ions into the serpentine structure is crystallographically favorable because of
the smaller ionic radius of Fe3+ compared to that of Mg2+ [53]. Moreover, based on a thermodynamic
model for hydrothermal alteration in the Fe-silicate system, it has been shown that the formation ofserpentine–hisingerite solid solutions after primary olivine may occur at elevated aSiO2(aq) and low
aH2(aq) at low temperatures (about 200 ◦ C) [28]. In addition, it has been suggested that H2 production
is associated with Fe(III) incorporation into serpentine (or magnetite) [54–56].
The association of magnetite with sulphides from the Pindos (Tsoumes) resembles an unusual
association of magnetite with sulphides of Cyprus-type ophiolite-hosted VMS deposit in Ortaklar,
located in the Koçali Complex, Turkey [7]. They are similar in terms of the deposition order, as suggestedby the observed textural relationships: Fe-sulphide (pyrrhotite or pyrite) → chalcopyrite → sphaleriteand subsequently magnetite (Figure 4; [7]) and the Cu, Zn, Pb, and Au-Ag contents [7]. Such temporal
relationships among the primary ore minerals have been attributed to the evolution of ore-forming
ﬂuids. Speciﬁcally, increasing oxygen fugacity (f O2 ) and pH would deplete sulphide (H2 S) and facilitate
the magnetite precipitation in the hydrothermal ﬂuids [7]. In addition, the Pindos sulphides exhibit
similarities with the Cogne magnetite deposit (Western Alps, Italy), which is the largest in a series of
apatite and sulphide-free magnetite orebodies that are hosted in serpentinites belonging to western
Alpine ophiolitic units [8]. The authors applying thermodynamic modelling of ﬂuid–rock interactions
concluded that fractionation processes such as phase separation were critical to generate hydrothermal
ﬂuids capable of precipitating large amounts of magnetite in various types of ultramaﬁc host rocks [8].
Although variable textures described in the large Cogne magnetite deposit diﬀer from those in the
Pindos magnetite ore, the trace element content of magnetite from the Cogne deposit, characterizedby high Mg and Mn and low Cr, Ti, and V [8], is comparable to those in the Pindos magnetite
separates [57]. Additionally, Fe–Cu–Zn–Co–Ni mineralization has been reported in seaﬂoor VMS
deposits from modern oceans, as well as in their potential analogues on several ophiolite complexes,
as exempliﬁed in the Urals, supporting the origin of such deposits by hydrothermal processes [19].
Moreover, hydrothermal products including Cu–Zn–(Co)-rich massive sulphides, hosted in ultramaﬁcrocks at the Rainbow (Mid-Atlantic Ridge), exhibit structure, mineralogy, and bulk rock chemistry
similar to those found in maﬁc volcanic-hosted massive sulphide deposits [58].
In general, the main factors controlling metal associations in seaﬂoor massive sulphide (SMS)
deposits may be the temperature of deposition, seaﬂoor spreading rate and r/w ratio, and zone
reﬁning [59]. The authors emphasized the signiﬁcance of the ﬁnal depositional conditions andevolution of mound and vent structures rather than the original geochemistry of the hydrothermal
ﬂuid; the composition of the substrate may become relevant in subseaﬂoor mineralization, wheresulphides are precipitated by the reaction of ascending hydrothermal ﬂuids with substrate hostrocks [59]. Individual deposits may show a mixture of geochemical signatures, which may be related
to maﬁc and ultramaﬁc rocks [19,58]. Assuming that the leaching of elements from substrate rocks is
inﬂuenced by the structure of the oceanic lithosphere and by the nature of the hydrothermal convection(spreading rate), some speciﬁc geochemical features, such as Au enrichment, Au/Ag and Co/Ni ratios,
may be related to the nature of the substrate, the presence of a magmatic inﬂux of volatiles and metals,
the morphology of vent structures, the ridge spreading rate, or a combination of these factors [19,58,59].
Although the observed textural and mineralogical features (Figures 3–5) are inconsistent with an originof the sulphides at magmatic temperatures, the recorded diﬀerences, such as the higher Zn, Se, Mo,
Au, Ag, Hg, and Sb and lower Ni contents in the Pindos compared to the Othrys sulphides (Table 3;
Figure 6), may reﬂect inheritance of a primary magmatic signature.
6.2. Genetic Signiﬁcance of Trace Elements
The massive Cyprus-type sulphides from the Kondro Hill are characterized by elevated Zn, Co,
Se, Au, As, Ag, Mo, and Sb content (Table 1) compared to those of the Othrys and most of the Troodos
sulphide ores (Figure 6). Apart from the major elements, namely Fe, Cu, and Zn, hosted in pyrite,

232

Minerals 2020, 10, 454

chalcopyrite, bornite, and sphalerite, Au occurs as submicroscopic particles (<1 μm) in grains ofAs-bearing pyrite, chalcopyrite, and bornite [20,26]. Selenium and Te are found as individual ﬁne
minerals, such as selenides (clausthalite) and tellurides (melonite) in Cu minerals (Figure 2f; [26]).
Although Se can be easily hosted as a solid solution in high-temperature chalcopyrite [58], the presenceof clausthalite in a late generation of ﬁne-grained chalcopyrite–sphalerite intergrowths, penetrating
into an earlier stage ore, may indicate re-distribution of Se. A late growth of clausthalite is also
supported by the occurrence of clausthalite and tellurides ﬁlling cracks in pyrite and Cu minerals [26].
Furthermore, on the basis of thermodynamic calculations, it has been demonstrated that the presenceof selenides in the oxidation zones of sulphide ores of Uralian VMS deposits is related to theirstability under oxidizing conditions [60]. Molybdenite or other visible Mo minerals in the Pindos andOthrys sulphides have not yet been reported. It has been established that Mo displays siderophile,
chalcophile, and lithophile behavior, depending on the composition of the system (including fO2 and
fS2 ), temperature, and pressure [61]. Further research is required to deﬁne the potential presence ofinvisible Mo minerals (less than 1 μ) in the Pindos sulphide-magnetite ores.
6.3. Stability of Sulphides
A salient feature of the sulphide minerals is a varying stability. The occurrence of euhedral
pyrite crystals, in contrast to microcrystalline unhedral Cu and Zn sulphides (Figure 2), may indicatethat pyrite was more stable during subsequent modiﬁcation of the orebody. Although early large
crystals of pyrite may be replaced by chalcopyrite or bornite and Fe-poor sphalerite intergrowths(Figure 2), the formation of cruciform aggregates of secondary minerals occurs only on bornite surfaces(Figure 2i–l), probably reﬂecting a diﬀerence in their stability. The preferential leaching of Cu and Zn
sulphide phases and the neo-formation of high Cu sulphides on bornite in contrast to neighboring
pyrite may be the result of a preferred dissolution of Cu sulphides over pyrite, due to diﬀering surface
potentials [62]. It has been suggested that bornite with sulfur in excess (x-bornite) is stable at hightemperature [62]. The authors of this study show that if the so-called sulfur-rich bornites are annealed
at lower temperature, chalcopyrite or chalcopyrite and digenite exsolve, depending on the annealing
temperature and composition. In addition to this exsolution, a new phase forms below approximately
140 ◦ C, which is referred to as x-bornite, and it is a metastable phase. Although x-bornite is a metastablephase, the presented data (Figure 2i–l; Table 1) may conﬁrm that x-bornite can remain for a longtime in natural environments, and epigenetic minerals, with a stoichiometry resembling geerite orspionkopite [63,64], can be formed under environmental conditions in a short time.
7. Conclusions
The compilation of the mineralogical, geochemical, and mineral chemistry data from the sulphideoccurrences hosted in the Pindos ophiolite complex and those from other ophiolites lead us to thefollowing conclusions:
•
•

•

•

Elevated contents of Au as invisible submicroscopic Au in pyrite and Cu minerals in the Pindos
sulphides may reﬂect main collectors of Au at the time of the sulphide mineralization.
The occurrence of clausthalite (PbSe) and ﬁne-grained gold in chalcopytite–bornite–sphalerite
intergrowths of a subsequent stage mineralization in the Pindos sulphides indicates theirre-mobilization/re-deposition.
Sulphides (pyrrhotite, chalcopyrite, bornite, and sphalerite) associated with magnetite, at deeper
parts of the Pindos (Tsoumes), exhibit Cu/(Cu + Ni), Ni/Co, and Pt/(Pt + Pd) ratios, suggesting
either no magmatic origin or a complete transformation of a preexisting magmatic assemblages.
Textural features and the presence of the (Fe/Mg) phyllosilicate resembling Mg–hisingerite,and calcite in the matrix of the Pindos sulphides, suggest precipitation of the sulphide-magnetite
ore at the deeper levels from a Fe-rich and alkaline ore-forming system.

233

Minerals 2020, 10, 454••

The preferential leaching of Fe and S and neo-formed high Cu sulphides on bornite, in contrast
to neighboring pyrite, may be the result of a preferred dissolution of Cu sulphides over pyrite,
conﬁrming literature data on diﬀering surface potentials between those sulphides.
Assuming that trace elements in epigenetic minerals are derived from the decomposition ofprimary minerals, and coupled with the higher Zn, Se, Mo, Au, Ag, Hg, and Sb and lower Nicontents in the Pindos compared to the Othrys sulphides, this may reﬂect inheritance of a primarymagmatic signature.

Author Contributions: D.G.E., M.E.-E., and G.E. collected the samples, provided the ﬁeld information,and contributed to the conceptualization of the manuscript. V.S. performed the SEM/EDS analyses.
M.E.-E. discussed the mineralogical and chemical data with the co-authors and carried out the original draft of themanuscript. All authors have read and agreed to the published version of the manuscript.
Funding: The National and Kapodistrian University of Athens (NKUA) is greatly acknowledged for the ﬁnancial
support (Grant No. KE_11730) of this work.
Acknowledgments: We thank the reviewers for the constructive criticism and suggestions on an earlier draft ofthe manuscript. In particular, the review of this work by the Academic Editor Paolo Nimis is greatly appreciated.
Many thanks are due to our colleague Costas Mparlas for the donation of certain sulphide samples from hiscollection and valuable discussions.
Conﬂicts of Interest: The authors declare no conﬂict of interest.

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minerals
ArticleA Comparative Study of Porphyry-Type Copper
Deposit Mineralogies by Portable X-ray Fluorescenceand Optical Petrography
Connor A. Gray and Adrian D. Van Rythoven *Department of Environmental, Geographical, and Geological Sciences, Bloomsburg University of Pennsylvania,
Bloomsburg, PA 17815, USA; connorgray@mymail.mines.edu* Correspondence: avanrythov@bloomu.edu; Tel.: +1-570-389-3912
Received: 30 March 2020; Accepted: 8 May 2020; Published: 11 May 2020

Abstract: Porphyry-type deposits are crucial reserves of Cu and Mo. They are associated withlarge haloes of hydrothermal alteration that host particular mineral assemblages. Portable X-ray
ﬂuorescence analysis (pXRF) is an increasingly common tool used by mineral prospectors to makejudgments in the ﬁeld during mapping or core logging. A total of 31 samples from 13 porphyry copper
deposits of the Western Cordillera were examined. Whole-rock composition was estimated over threepoints of analysis by pXRF. This approach attempts to capture the rapid and sometimes haphazard
application of pXRF in mineral exploration. Modes determined by optical petrography were converted
into bulk rock compositions and compared with those determined by pXRF. The elements S, Si,
Ca, and K all were underestimated by optical mineralogy, and the elements Cu, Mo, Al, Fe, Mg,
and Ti were overestimated by optical mineralogy when compared with pXRF results. Most of theseporphyry samples occur in veined porphyritic quartz monzonite that is characteristic of these deposits.
Sulﬁde and silicate vein stockworks are pervasive in most of the samples as well as dissemination of
sulﬁdes outwards from veinlets. Ore minerals present include chalcopyrite and molybdenite withlesser bornite. Chalcocite, digenite, and covellite are secondary. Potential sources of analytical bias
are discussed.
Keywords: Western Cordillera; pXRF; North America; modal mineralogy; optical mineralogy;
porphyry deposit; copper; molybdenum; major element; handheld X-ray ﬂuorescence

1. Introduction
1.1. Porphyry-Type Deposits
Porphyry-type deposits, often termed just “porphyry deposits”, represent one of the most studied
geological systems on Earth, and the economic fortunes that entail from the proper extraction of
the metals found in these deposits have pushed the academic and industrial communities to study
them. This deposit model that came to popularity in the mid-late 1960s [1,2], was summarized by
Sillitoe [3] and then later expanded upon [4–8]. Porphyry-type deposits represent 60% of world
Cu production annually, and 65% of known world Cu resources [6]. These deposits also represent
half of the world’s Mo production and may also contain economic grades of Ag, Zn, W, Sn, Re,
and Au [7]. Low-grade resources of Pt-group elements are also possible [9,10]. Tonnages commonly
exceed one billion [8]. An example of a deposit of this scale is Bingham Canyon (Utah), a deposit
analyzed in this study, at 2.6 Gt of ore [11]. These high tonnages are oﬀset by lower grades of metals
extracted, typically between 0.5–1.0 wt% Cu [7]. The deposits are typically found along convergent
plate margins where the subduction of an oceanic plate below the adjacent continental or oceanic platedrives partial melting, thus creating arc magmatism [3]. Although similar types of deposits occur inMinerals 2020, 10, 431; doi:10.3390/min10050431

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rarer continental-continental orogens [12]. These magmatic arc systems allow metals found in greater
concentrations deeper in the earth to rise buoyantly as magma and emplace themselves as intrusive
bodies capable of metal concentration. These systems cool, and depressurization processes occur suchthat metals are highly concentrated in an immiscible hydrothermal ﬂuid located at the top of such
intrusive bodies. These metals are then injected into the surrounding rock by hydraulic fracturingforming large stockworks of cross-cutting veins that contain concentrated metals [8]. The particularsamples analyzed in this study are from the Western Cordillera of the Americas. Deposits of thistype are essential to the growth of economies in developing countries, such as Chile where copper
mining has contributed an average of 10% of GDP over the past 20 years [13]. Deposits found along
the Western Cordillera of North America are, or have been inﬂuential regarding the economies of theregions and countries these deposits are found in.
1.2. Portable X-ray Fluorescence
X-ray ﬂuorescence analysis produces compositional data by exciting the sample in question with
X-rays. Those X-rays will ionize the sample’s atoms by knocking out core-shell electrons. Outer-shellelectrons will then fall/relax into the subsequent core-shell vacancy. As this involves a loss of energy,
a photon is emitted with energy equal to the diﬀerence in electron orbitals. The energy quantum of thephoton is unique to the element in question and is also in the X-ray portion of the electromagnetic
spectrum. Thus, the X-ray’s energy is characteristic of the element and the number of photons withthat quantum is proportional to the element’s concentration in the sample. Engineering advances
over the last few decades have allowed this laboratory-scale instrumentation to be miniaturized into
a portable format. A variety of commercial portable and handheld X-ray ﬂuorescence analyzers
are now available [14]. Portable X-ray ﬂuorescence (pXRF) is an increasingly common method foron-site and rapid materials analysis. For exploration and development of mineral deposits, it has
particularly been applied to geochemical exploration [15–17], assessments of soil contamination [18],core logging [19], and mineralization type [20,21]. This is a method that requires little training,
is portable, fast, fairly aﬀordable, and can provide useful data. However, compared to other lab
methods (including more sophisticated, lab-based, and non-portable XRF analyzers), analyses canbe very susceptible to measurement bias, insensitivity to elements lighter than Mg or Al, and matrix
eﬀects [22,23]. Whereas “portable” can refer to a variety of XRF conﬁgurations, for this study it refers
to handheld instruments, typically resembling a pistol.
As pXRF is often used as a tool by ﬁeld or core logging geologists, it is typically one of theﬁrst analytical methods applied to samples. The absence of typical sample preparation in the ﬁeld,
the small sample point size, and human bias in selecting what points to analyze introduce additional
uncertainties to the data. Although they may be collected under less than optimal analytical conditions,
results from pXRF can inﬂuence decisions regarding geologic mapping, cross-sections, and selection ofsamples for more detailed assay/investigation. Another investigative method that can follow pXRF
is optical petrography. In fact, data from pXRF may inﬂuence a geologist to select a sample for thinsectioning and petrographic examination. A geologist can later re-interpret and reconcile the pXRF datain the context of sample mineralogy. Likewise, pXRF data can be useful in interpreting the observed
mineralogy in thin sections. The combined use of complementary but independent methods such as
pXRF and optical petrography has the potential to augment the abilities of the investigating geologist
in a fairly fast and economical manner.
In addition to its portability and ease of use, pXRF is typically non-destructive. Outcrop, drill core,grab samples, soil, and thin section billets (this study) can be analyzed with no crushing, grinding,fusion, dissolution, etc., that lab-based XRF or ICP-MS require. Similarly, polished thin sections that canbe made from a modest amount of sample, are not expended when examined, preserve mineral textures,and can later be analyzed by other methods such as electron microscopy or other microbeam methods.
We present that pXRF used in a rapid fashion with no standardization, in the manner of many ﬁeld
applications, can still provide semi-quantitative data that are useful to answer geological questions in240

Minerals 2020, 10, 431

the ﬁeld. We simulate these haphazard conditions using point analysis on thin-section billets of rocksamples from porphyry-type deposits. Sources of analytical uncertainty arising from these conditions
are examined and discussed. The results are compared with those of optical petrography, anothercommon, albeit much older, standby in non-destructive geological sample characterization.
2. Materials and Methods
2.1. Sample Locations
The 31 hand samples used in this comparative study are from 13 diﬀerent deposits found along
the oceanic-continental subduction zone of the Western Cordillera of North America (Table 1, Figure 1).
These are Cu ± Mo porphyry-type deposits and represent diﬀerent portions within each deposit(for those deposits with more than one sample). The samples include specimens from the hypogene
and supergene zones of these deposits. The zone of interest for this study was primarily the hypogene.
These samples were collected and donated by Peter H. Kirwin and John R. Ray, formerly of theAmerican Copper and Nickel Company.
Table 1. List of samples investigated in this study and their origins.
SampleLocality

Location

SampleLocality

Location

Be-1
Be-4
Be-6
Bi-2
Bi-3
Bu-2
Bu-6
Bu-7
C-3
C-4
CC-1
F-1

Bethlehem
Bethlehem
Bethlehem
Bingham Canyon
Bingham Canyon
Butte
Butte
Butte
Cananea
Cananea
Copper Canyon
Florence

British Columbia
British Columbia
British Columbia
Utah
Utah
Montana
Montana
Montana
MexicoMexicoNevadaArizona

L-1
Mo-2
Mo-4
Mo-6
Mo-8
OH-1
S-1
S-2
S-4
SM-3
SM-4
Y-3

LornexMorenciMorenciMorenciMorenciOrange Hill
Sierrita
Sierrita
Sierrita
San Manuel
San ManuelYerington

British Columbia
Arizona
Arizona
Arizona
Arizona
AlaskaArizona
Arizona
Arizona
Arizona
Arizona
Nevada

F-2
F-3
IP-1
IP-2
IP-3
IP-10
IP-12

Florence
Florence
Ithaca Peak
Ithaca Peak
Ithaca Peak
Ithaca Peak
Ithaca PeakArizona
Arizona
Arizona
Arizona
Arizona
Arizona
Arizona

2.2. Analytical Methods
In this comparative study two methods were used to determine the overall mineralogy of thesamples. These methods include using pXRF giving elemental weight percent readings for each sample,
as well as using polarized transmitted and reﬂected light microscopy to estimate mineral abundancesand related textures. Limiting the study to these two methods allows the investigation to be conductedin an essentially non-destructive manner at minimal cost.

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Minerals 2020, 10, 431

D

Figure 1. Cont.

242

Minerals 2020, 10, 431

FFigure 1. Cont.

243

Minerals 2020, 10, 431

G
Figure 1. Maps of deposit locations in western North America that were sampled. (a) Overall geologic
map of the Western Cordillera (map constructed in ArcGIS). (b) Satellite image of the Orange Hill
deposit in Alaska. (c) Satellite image of the British Columbian, Nevadan, Montanan, and Utahan
deposits. (d) Satellite image of Arizonan and Mexican deposits.

2.2.1. Optical Petrography
The 31 samples were selected from a larger collection on the basis of interesting mineralogy or
textures visible in hand samples, geographic location, and budget limitations (for thin sectioning).
They were cut into rectangular billets by a rock saw (e.g., Figure 2). The billets were then used toshave oﬀ 30 μm × 26 mm × 46 mm polished thin sections for further analysis under a petrographic
microscope. The thin sectioned samples were then examined under a Leica DM2700 P microscope(Leica Microsystems Inc., Buﬀalo Grove, IL, USA) in transmitted and reﬂected light modes. In the ﬁrst
assessment of these thin sections, the textures were documented along with a qualitative listing ofminerals present in each sample. More in-depth qualitative observations were then made at this point,
and the genetic sequencings of minerals along with their characteristics were noted. In the second
extensive assessment of these samples using petrographic methods, the overall mineral modes were
determined. This assumed that the mineral abundances in terms of thin section area are representativeof the volume% abundances (true modes) of the mineral.

244

Minerals 2020, 10, 431

D

E

G

F

H

I

Figure 2. Example photographs of two samples from this study. (a) Bu-2 (Butte, Montana) handsample showing sulﬁde vein stockwork characteristic of porphyry-type deposits and alteration haloes
surrounding the veins. (b) Thin section billet cut from the sample in (a) and impregnated with epoxy.
(c) Be-6 (Bethlehem, British Columbia) hand sample showing sulﬁde vein stockwork characteristicof porphyry-type deposits and alteration adjacent to the veins. (d) Thin section billet cut from thesample in (c) and impregnated with epoxy. (e) IP-12 (Ithaca Peak, Utah) hand sample showing ahigh-grade vein of Cu-Fe sulﬁdes. (f) Y-3 (Yerington, Nevada) somewhat brecciated hand sample
showing cross-cutting veinlet stockwork formed in multiple generations of fractures. American quarterfor scale (~2.43 cm in diameter). Locations to be cut for thin-sectioning are outlined in black marker.

245

Minerals 2020, 10, 431

2.2.2. Portable X-ray Fluorescence
When the overall modes were determined for each sample, the bulk elemental compositions werethen determined by pXRF analysis of the ﬂat face of the corresponding billet the thin section was cut
from. A Thermo Scientiﬁc NitonTM XL3t-950 GOLDD+ (Thermo Fisher Scientiﬁc, Waltham, MA, USA)was the pXRF analyzer used in this study. All analyses were done in Mining Cu/Zn mode. A 10-minute
warmup and system check were conducted prior to every session of analyses. Analysis quality waschecked for instrument drift using two manufacturer-provided standards: CRM 180-649 NIST 2709a
and blank 180-647 99.995% SiO2 blank that were checked using the pXRF prior to the study and after
the study. These standards are supplied as ﬁne homogenized powders. No signiﬁcant deviation fromknown values (>10%) was noted. Samples were analyzed with pXRF for a 120-s interval, with 30 s eachfor the main (Mo, Zr, U, Rb, Th, Pb, Se, As, Hg, Nb, Cu, Ni, Co, Fe, Mn, Sr, Y), low (Sc, Ca, K, V, Ti, Cr),high (Ba, Cs, W, Te, Sb, Sn, Cd, Ag, Pd, La, Ce, Pr, Nd), and light (Al, P, Si, Cl, S, Mg) characteristic X-ray
ranges of elements using a series of ﬁlters. This allowed for fairly precise readings of elemental weightpercent (wt%) for each sample. Each sample was analyzed three times in three separate locations
on the sectioned surface of the billet in an attempt to address inhomogeneities within these billets.
The average elemental composition was determined using all three readings per sample. In the case of
three very diﬀerent analyses, a fourth location was analyzed on the billet.
2.3. Data Calculation
Modes from petrographic microscopy observations were converted into estimates of bulk elemental
composition for each sample. The modes, taken as volumetric percentages, were converted into mineralmass percentages using established densities. These mineral mass proportions were then converted
into elemental proportions. Density and compositional values were sourced from webmineral.com.
Elements not analyzable by pXRF (mostly elements lighter than Mg) were left out of calculations in themodal abundance conversion to allow for simpler comparison. Estimates of solid solution members
were made for plagioclase, chlorite, and muscovite. In terms of atoms per formula unit (apfu) it isassumed that plagioclase has 0.5 Na, 0.5 Ca, 1 Al, and 3 Si, chlorite has 3.75 Mg and 1.25 Fe, and biotite
has 2.5 Mg and 0.5 Fe.
The whole-rock compositions calculated from the modes were then compared with the averagewhole-rock compositions determined by pXRF for purposes of comparison and to audit the accuracyof each method against the other.
3. Results
3.1. Hand Sample Examination
At the hand sample scale, prior to the thin sectioning of each sample, porphyry-type copper deposit
characteristics were evident (e.g., Figure 2; see [24] for a review). Characteristic veining associated withhydraulic fracturing was seen in many samples as mineral assemblages of silicates and sulﬁdes withrare carbonates. Sulﬁdes associated with the hypogene system are typically dominated by chalcopyrite
and pyrite. Silicates in this zone displayed less alteration than that of more near-surface samples,maintaining the dominantly quartz monzonitic mineralogy seen in a majority of hypogene samples
although chalky sericitization of feldspars is common in the host rocks. In hand samples, characteristic
zoning of alteration could be seen with some ambiguity but overall a grade from hypogene to supergene
alteration was clear (c.f., [7]). The oxidation and hydration of hypogene sulﬁde minerals was a clear
indicator of supergene alteration.
3.2. Optical Petrography
In thin sections, the observations made of the hand samples were expressed in clearer detail, and theoverall modes observed in hand samples are consistent with those observed at a microscopic level(Table 2, Figure 3). The main sulﬁdes present in hypogene samples show characteristic veins/veinlets
246

Minerals 2020, 10, 431

of chalcopyrite, bornite, and/or pyrite (e.g., Figure 3g,f), along with sulﬁde grains disseminated distallyfrom veinlets (e.g., Figure 3a,h). These veinlets also comprise hydrothermal quartz along with sericite
and chlorite, and rarely calcite (e.g., Figure 3g). Molybdenite was the only Mo-bearing sulﬁde seen inthin section (e.g., Figure 3b,c,h, Table 2). Samples aﬀected by some supergene ﬂuid alteration havedistinct mineralogies. The sulﬁde assemblages of the hypogene zone, though starting rich in Cu
sulﬁdes, quickly grade into zones of pyrite showing a lack of Cu concentration distally (e.g., Figure 3f).
In the supergene samples, chalcocite, covellite, and/or digenite are observed as rims on partiallyreplaced pyrite and/or chalcopyrite (Figure 3a). Specular hematite (possibly with magnetite in somecases) occurs in some samples and also appears to be hydrothermal. This is distinct from the latehematite after sulﬁdes seen in some partially oxidized samples (see below).
Table 2. Mineral modes (in volume%) of samples. Mineral abbreviations as in Whitney and Evans [25].
Sample:

Be-1

Be-4

Be-6

Bi-2

Bi-3

Bu-2

Bu-6

Bu-7

C-3

C-4

CC-1

F-1

F-2

F-3

IP-1

IP-2

Ccp
Cv
Cct
Bn
Dg
Mol
Py
Chl
Rt
Hem
Cal
Qz
Or
Ms
Bt
Pl
Ap
Total

0.3
–
–
–
–
–
1
8
0.3
2
–
57.4
–
27
–
3
1
100.0

1
–
–
1
–
–
–
4.5
–
0.1
2
64.9
–
25.5
0.5
–
0.5
100.0

1.5
–
–
0.2
–
–
–
7
0.3
–
–
60
–
22
–
8
1
100.0

1
0.3
1.5
0.1
–
0.2
2
1
0.5
–
–
48.7
–
42.7
1
–
1
100.0

1
–
0.1
1
0.3
0.1
–
–
0.2
–
–
61.3
–
25
10
–
1
100.0

0.5
0.2
–
–
–
–
9
–
1
3
–
56.3
–
30
–
–
–
100.0

–
–
0.5
–
–
–
20
–
–
–
–
74.5
3
2
–
–
–
100.0

1.5
0.3
–
–
–
0.5
45
–
–
–
–
48.7

–
–
–
–
–
0.5
4
–
–
–
–
60.5

–
2
–
–
–
7
–
–
0.5
–
50

1
–
–
–
–
–
–
–
0.3
–
–
67

4
–
–
–
–
–
–
–
0.5
0.5
–
58

4
–
–
–
100.0

35
–
–
–
100.0

40.5
–
–
–
100.0

23.7
8
–
–
100.0

34
3
–
–
100.0

0.8
–
–
–
–
–
2
0.4
1
0.5
–
56.3
10
12
10
5
2
100.0

2.5
–
–
–
–
–
2
0.3
0.4
–
3
59.8
–
32
–
–
–
100.0

–
–
0.5
–
–
–
3
–
1
–
–
64.5
–
31
–
–
–
100.0

0.1
1
0.5
–
–
–
4
–
–
–
–
67.4
–
27
–
–
–
100.0

Sample:

IP-3

IP-10

IP-12

L-1

Mo-2

Mo-4

Mo-6

Mo-8

OH-1

S-1

S-2

S-4

SM-3

SM-4

Y-3

Ccp
Cv
Cct
Bn
Dg
Mol
Py
Chl
Rt
Hem
Cal
Qz
Or
Ms
Bt
Pl
Ap
Total

0.1
0.1
–
–
–
0.3
1
–
1
1
–
43.3
16
10
24.5
1
1.7
100.0

0.05
0.1
0.2
–
–
–
9
–
–
–
–
85.65
–
5
–
–
–
100.0

22
–
–
10
–
–
–
5
–
0.5
–
44.5
1
12
1
4
–
100.0

6
–
–
–
–
0.1
–
–
–
–
–
85.4
–
8
–
–
0.5
100.0

1
–
–
–
–
–
5
13
1
–
–
45
–
34
–
–
1
100.0

–
–
–
–
–
–
2.5
–
0.3
–
–
94.9
–
2.3
–
–
–
100.0

2
–
–
–
–
–
–
1.5
0.3
–
0.3
54.4
–
34.5
2
5
–
100.0

1.5
–
–
–
–
–
1.4
–
0.2
–
0.9
53.4
7.1
33
2.5
–
–
100.0

1
–
–
–
–
–
8
11
–
0.5
–
26.5
–
15
3
35
–
100.0

1
–
–
–
–
–
1.6
11
0.5
–
–
35
8.9
10
–
32
–
100.0

2
–
–
–
–
0.2
2
2
0.5
–
–
62.5
4
25
–
1.3
0.5
100.0

0.5
–
–
–
–
0.3
3
0.5
0.5
0.3
1
50.9
3
36
–
4
–
100.0

3.5
–
–
–
–
–
5
–
0.5
–
–
43
–
37
4
6.3
0.7
100.0

4.5
–
–
–
–
–
1
–
1.5
1
–
36.3
–
26
25
4
0.7
100.0

2
–
–
–
–
–
–
4
0.3
1.5
–
55.7
5
25
–
6
0.5
100.0

The host rocks vary depending on the particular deposit. Most are porphyritic hypabyssal/intrusive
rocks that are quartz monzonitic/latitic, although granitic/rhyolitic and granodioritic/dacitic assemblages
occur (c.f., [26]). Primary quartz is present in varying amounts but siliciﬁcation is prevalent. Biotite,
apatite, rare amphibole, and rare zircon are also present as primary minerals. In addition to siliciﬁcation,
potassic alteration, sericitization (mostly muscovite after plagioclase and orthoclase) and chloritization(of biotite) are prevalent (e.g., Figure 3e). Carbonation (of plagioclase and maﬁc minerals) occurs but
is rare. In many samples, siliciﬁcation and/or sericitization obliterated much of the primary phases
and textures.
Supergene oxidation is evident in some samples by partial replacement of pyrite by masses ofhematite ± goethite. These samples also showed some spots of decomposition of Cu-bearing sulﬁdesas a greenish tinge or patina (i.e., chrysocolla-malachite) on hand samples, however, none of the areas
thin sectioned contained these spots to any signiﬁcance.

247

Minerals 2020, 10, 431

D

E

F

G

H

I

Figure 3. Cont.

248

Minerals 2020, 10, 431

J

K

Figure 3. Example photomicrographs of two samples from this study. (a) Bi-2, chalcopyrite withalteration to chalcocite/covellite along crystal rims and fractures. Surrounding phases are quartz
and sericite. RPPL (reﬂected plane-polarized light). (b) Bi-3, chalcopyrite, bornite, and molybdenite
intergrown. Digenite alteration along the bornite and chalcopyrite crystal rims. Surrounding phases
are biotite, quartz, and sericite. RPPL. (c) S-2, chalcopyrite and molybdenite in a quartz vein. RPPL(d) Bu-6, large pyrite crystals (some with possible zircon inclusions) in quartz + sericite + orthoclase.
TXPL (transmitted cross-polarized light). (e) IP-12, vein of chalcopyrite + bornite + biotite in sericitizedand chloritized porphyritic biotite quartz monzonite/latite. TPPL (transmitted plane-polarized light).
(f) Same view as in (e), but in RPPL. (g) F-3, quartz + pyrite + chalcopyrite + calcite vein in sericitizedmonzonite host. (h) L-1, small grains of chalcopyrite (lower left) and molybdenite (upper right) inquartz and sericite. RPPL.

3.3. Estimation of Bulk Compositions from Modes
The estimated bulk rock compositions for these modes was calculated by converting the observed
mineral abundances to weight proportions, and then elemental proportions based on idealizedcompositions of the minerals. Based on the observed minerals, only the elements Mg, Al, Si, P, S, K, Ca,
Ti, Fe, Cu, and Mo were considered (Table 3).
Table 3. Bulk rock compositions (elemental wt%) for selected elements calculated from modes
determined by optical petrography.
Sample

Be-1

Be-4

Be-6

Bi-2

Bi-3

Bu-2

Bu-6

Bu-7

C-3

C-4

CC-1

F-1

F-2

F-3

IP-1

IP-2

Mg
Al
Si
P
S
K
Ca
Ti
Fe
Cu
Mo
Balance

1.2
6.3
32.3
0.2
1.1
2.7
0.5
0.3
1.9
0.2
0.0
53.5

0.8
5.8
35.5
0.1
1.0
2.6
1.0
0.0
1.4
1.7
0.0
50.1

1.0
6.1
35.4
0.2
0.9
2.2
1.0
0.3
1.6
1.0
0.0
50.2

0.3
8.7
30.2
0.2
3.4
4.2
0.4
0.4
2.3
3.3
0.2
46.2

1.6
5.8
34.5
0.2
1.2
3.5
0.5
0.2
1.4
2.3
0.1
48.7

0.0
5.7
28.9
0.0
8.4
2.8
0.0
0.8
10.9
0.4
0.0
42.0

0.0
0.6
30.4
0.0
17.4
0.5
0.0
0.0
15.0
0.7
0.0
35.4

0.0
0.6
16.5
0.0
33.2
0.3
0.0
0.0
28.5
0.8
0.4
19.6

0.0
7.2
33.9
0.0
4.2
3.5
0.0
0.0
3.3
0.0
0.6
47.3

0.0
7.9
29.2
0.0
7.5
3.8
0.0
0.0
6.2
2.1
0.0
43.2

1.3
5.5
37.0
0.0
0.5
3.2
0.0
0.3
1.1
0.5
0.0
50.5

0.5
7.2
33.4
0.0
2.1
3.7
0.0
0.5
2.7
2.1
0.0
47.9

1.6
4.6
33.8
0.4
2.3
3.5
1.3
0.9
3.5
0.4
0.0
47.8

0.0
6.6
33.2
0.0
3.2
3.2
1.2
0.4
3.1
1.3
0.0
47.7

0.0
6.4
35.0
0.0
3.1
3.1
0.0
0.9
2.5
0.8
0.0
48.2

0.0
5.5
35.1
0.0
4.6
2.7
0.0
0.0
3.4
2.0
0.0
46.7

Sample

IP-3

IP-10

IP-12

L-1

Mo-2

Mo-4

Mo-6

Mo-8

OH-1

S-1

S-2

S-4

SM-3

SM-4

Y-3

Mg
Al
Si
P
S
K
Ca
Ti
Fe
Cu
Mo
Balance

3.7
5.2
30.7
0.3
1.3
5.4
0.8
0.9
3.9
0.2
0.3
47.2

0.0
1.0
37.8
0.0
8.6
0.5
0.0
0.0
7.4
0.5
0.0
44.3

0.8
0.3
20.9
0.0
13.9
1.2
0.2
0.0
11.5
19.7
0.0
31.5

0.0
1.7
40.0
0.1
3.3
0.8
0.2
0.0
2.8
3.2
0.1
47.8

1.9
7.9
28.2
0.2
5.2
3.3
0.4
0.9
6.0
0.5
0.0
45.5

0.0
0.5
43.7
0.0
2.5
0.2
0.0
0.3
2.2
0.0
0.0
50.6

0.5
8.0
34.0
0.0
1.1
3.7
0.5
0.3
1.2
1.1
0.0
49.7

0.4
7.7
33.4
0.0
2.2
4.5
0.4
0.2
2.1
0.8
0.0
48.5

2.0
7.7
26.3
0.0
7.9
1.7
2.4
0.0
8.9
0.5
0.0
42.7

1.6
7.0
31.8
0.0
2.1
2.2
2.3
0.5
3.1
0.5
0.0
48.8

0.3
5.8
34.9
0.1
3.2
3.0
0.3
0.5
2.8
1.0
0.2
47.8

0.1
8.1
32.0
0.0
3.4
3.9
0.7
0.5
3.2
0.3
0.4
47.6

0.6
8.1
28.4
0.1
6.4
3.9
0.7
0.4
5.8
1.7
0.0
43.7

3.7
7.1
26.6
0.1
3.2
4.8
0.6
1.3
5.7
2.2
0.0
44.8

0.6
6.6
34.0
0.1
1.1
3.2
0.7
0.3
3.4
1.1
0.0
49.1

249

Minerals 2020, 10, 431

3.4. Portable X-ray Fluorescence Analysis of Bulk Compositions
Analysis by pXRF provided compositional information in terms of elements Mg and heavier(Table 4). Overall composition across the samples varies signiﬁcantly. Signiﬁcant variations in Mg, Si,
P, S, K, Ca, Ti, Fe, Cu, and Mo content occur. Other elements analyzed for, but are low in concentration(<0.5 wt%) and do not show signiﬁcant variation are: Cl, V, Cr, Mn, Co, Ni, Zn, As, Se, Rb, Sr, Y, Zr, Nb,
Ag, Cd, Sn, Sb, Ba, La, Ce, Pr, Nd, W, Au, Pb, Th, and U. Concentrations for V, Co, Se, Y, Sn, Sb, Au,
and U were essentially non-detectable (less than approximately 0.001 wt%). Uncertainty in analyses isdominated by the variance between spots compared to the actual instrumental error with the possible
exception of Mg, the element with the highest uncertainty in terms of measurement.
In terms of the economic metals, Cu grades vary from 0 to 14.8 wt% (mean: 0.8 wt%) and Mo
grades vary from 0 to 0.3 wt% (mean 0.02 wt%). For the other elements, silicon is the most abundant
and ranges from 23.2 to 53.8 wt% (mean: 37.4 wt%). Al is the second most consistently abundantelement (0.1 to 8 wt%, mean: 4.9 wt%) followed by K (0 to 8.5 wt%, mean: 3.5 wt%). Elements that arehigh abundance in some samples (those with high sulﬁde mineral content) are Fe (0.5 to 20.8 wt%,
mean: 3.6 wt%) and S (0.1 to 40 wt%, mean: 5.4 wt%). Less signiﬁcant variations are Mg (0 to 3.5 wt%,
mean: 0.6 wt%), P (0 to 0.4 wt%, mean 0.1 wt%), Ca (0 to 3.2 wt%, mean 0.8 wt%), and Ti (0 to 1.2
wt%, mean 0.3 wt%). The sum of chalcophile element concentrations: Fe, Cu, and Mo, correlate with S
content (Figure 4).

&ĞнƵнDŽ;ǁƚйͿďǇƉyZ&

ϯϬ
Ϯϱ
ϮϬ
ϭϱ
ϭϬ
ϱ
Ϭ
Ϭ

ϭϬ

ϮϬ

ϯϬ

ϰϬ

ϱϬ

^;ǁƚйͿďǇƉyZ&
Figure 4. Plot of S content vs. sum of chalcophile elements determined by pXRF.

3.5. Comparison of Bulk Compositions from Both Methods
Calculated bulk rock composition (in terms of selected elements) based on estimated modes fromoptical petrography and the estimated bulk rock compositions based on pXRF analyses are comparedon an element-by-element basis (Figure 5).

250

Be-4

0.422
0.201
0.081
5.798
0.136
2.363
41.352
0.207
7.025
0.110
0.023
0.018
0.581
0.014
0.625
3.145
0.039
1.162
1.282
0.029
0.276
0.113
0.059
0.017
1.012
0.024
0.158
0.442
0.006
0.632
0.000
0.002
0.000
45.743

Be-1

0.558
0.197
0.306
5.995
0.134
0.131
36.175
0.213
1.925
0.085
0.020
0.047
0.136
0.008
0.155
2.768
0.041
0.484
1.410
0.034
0.529
0.262
0.088
0.155
2.187
0.031
1.083
0.027
0.002
0.034
0.000
0.002
0.000
50.397

SampleMg
Mg(2σ)inst
Mg(s)spot
Al
Al(2σ)inst
Al(s)spotSi
Si(2σ)inst
Si(s)spotPP(2σ)inst
P(s)spot
S
S(2σ)inst
S(s)spot
K
K(2σ)inst
K(s)spot
Ca
Ca(2σ)inst
Ca(s)spot
Ti
Ti(2σ)inst
Ti(s)spotFeFe(2σ)inst
Fe(s)spot
Cu
Cu(2σ)inst
Cu(s)spotMo
Mo(2σ)inst
Mo(s)spotbalance

0.921
0.203
0.470
5.979
0.130
0.659
32.875
0.215
1.674
0.201
0.019
0.045
0.472
0.012
0.127
1.272
0.034
0.140
1.086
0.030
0.077
0.448
0.105
0.086
2.561
0.033
0.304
0.212
0.005
0.032
0.000
0.002
0.000
53.974

Be-6
0.610
0.284
0.107
6.977
0.173
0.077
33.531
0.201
1.190
0.372
0.026
0.100
2.778
0.036
0.470
8.525
0.083
0.160
0.137
0.026
0.029
0.414
0.115
0.026
1.187
0.027
0.210
1.658
0.016
0.513
0.005
0.001
0.002
43.805

Bi-2
0.776
0.237
0.088
6.007
0.147
0.646
37.622
0.209
0.584
0.150
0.023
0.079
0.564
0.015
0.187
7.306
0.069
0.781
0.231
0.023
0.088
0.228
0.083
0.146
0.760
0.024
0.169
0.445
0.007
0.116
0.000
0.002
0.000
45.910

Bi-3
0.332
0.668
0.423
5.821
0.185
0.348
34.213
0.208
7.114
0.071
0.035
0.057
12.5770.100
12.607
3.381
0.063
0.502
0.021
0.027
0.024
0.336
0.179
0.224
8.580
0.066
5.749
0.019
0.002
0.010
0.000
0.002
0.000
34.650

Bu-2
0.000
0.409
0.000
0.134
0.076
0.134
49.225
0.182
2.716
0.015
0.045
0.027
8.156
0.064
0.753
0.030
0.010
0.005
0.009
0.004
0.000
0.000
0.062
0.000
2.783
0.029
0.373
0.273
0.005
0.069
0.000
0.002
0.000
39.374

Bu-6
0.000
1.405
0.000
0.152
0.197
0.147
23.2430.137
20.339
0.000
0.070
0.000
39.9200.451
23.920
0.028
0.036
0.043
0.014
0.019
0.019
0.000
0.163
0.000
20.8060.422
15.386
0.414
0.014
0.262
0.002
0.002
0.003
15.422

Bu-7
0.000
0.318
0.000
6.972
0.153
0.954
39.127
0.220
2.561
0.000
0.033
0.000
1.933
0.028
0.798
4.740
0.055
0.493
0.054
0.016
0.006
0.429
0.094
0.027
1.311
0.026
0.483
0.064
0.003
0.019
0.000
0.002
0.000
45.370

C-3
0.328
0.356
0.294
6.819
0.163
0.641
36.880
0.198
2.991
0.020
0.033
0.034
6.521
0.060
2.363
3.737
0.051
0.246
0.080
0.017
0.010
0.201
0.092
0.035
3.047
0.035
0.597
0.186
0.005
0.045
0.000
0.002
0.000
42.181

C-4
0.559
0.209
0.247
5.109
0.131
0.215
40.978
0.217
1.034
0.000
0.039
0.000
0.056
0.007
0.006
5.275
0.056
0.651
0.460
0.024
0.135
0.087
0.083
0.083
1.207
0.026
0.116
0.309
0.005
0.061
0.000
0.002
0.000
45.960

CC-1
0.141
0.326
0.244
4.832
0.121
1.550
38.249
0.210
2.019
0.000
0.029
0.000
0.903
0.017
0.425
4.364
0.044
3.827
0.599
0.022
0.247
0.070
0.080
0.121
0.726
0.023
0.324
0.257
0.005
0.115
0.000
0.002
0.000
49.860

F-1
0.851
0.213
0.488
4.397
0.119
0.357
40.287
0.232
3.793
0.085
0.026
0.075
0.885
0.018
0.596
1.939
0.039
1.012
1.185
0.032
0.168
0.594
0.109
0.331
2.522
0.033
1.217
0.063
0.003
0.040
0.000
0.002
0.000
47.192

F-2
0.151
0.391
0.262
5.828
0.156
1.631
40.416
0.196
6.274
0.000
0.041
0.000
3.250
0.037
1.270
6.701
0.064
1.825
0.341
0.025
0.210
0.042
0.102
0.073
2.053
0.029
0.811
0.664
0.008
0.439
0.002
0.002
0.003
40.553

F-3
0.000
0.327
0.000
5.247
0.127
2.394
41.148
0.208
5.434
0.080
0.024
0.012
4.570
0.045
3.526
3.242
0.042
1.416
0.000
0.026
0.000
0.068
0.101
0.118
1.887
0.029
1.088
0.006
0.002
0.006
0.005
0.002
0.008
43.746

IP-1
0.000
0.437
0.000
4.555
0.123
0.411
36.948
0.187
5.187
0.000
0.038
0.000
6.029
0.052
5.011
3.886
0.044
0.881
0.028
0.014
0.007
0.034
0.082
0.058
1.657
0.027
1.582
0.674
0.007
1.114
0.043
0.001
0.042
46.146

IP-2

Table 4. Bulk rock compositions (elemental wt%) for selected elements determined by the average of point analyses for each thin section billet. Uncertainties are theaveraged instrumental population standard deviation determined by the pXRF software: (2σ)inst and the sample standard deviation of the spot analyses across thebillet: (s)spot . Samples with more than three spots analyzed are Bu-7 (ﬁve spots), Bu-2 (four spots), and IP-3 (four spots).

Minerals 2020, 10, 431

251IP-10

0.000
0.460
0.000
0.630
0.075
0.657
42.872
0.176
6.181
0.045
0.038
0.042
11.951
0.080
8.437
0.301
0.015
0.358
0.007
0.009
0.012
0.000
0.066
0.000
4.299
0.037
3.306
0.122
0.003
0.139
0.065
0.001
0.069
39.708

IP-3

3.466
0.349
2.952
3.792
0.132
2.270
33.8550.226
10.773
0.417
0.025
0.436
1.424
0.023
1.185
5.778
0.075
2.875
0.558
0.033
0.624
0.698
0.148
0.334
3.156
0.039
2.385
0.022
0.002
0.006
0.273
0.003
0.372
46.561

SampleMg
Mg(2σ)inst
Mg(s)spot
Al
Al(2σ)inst
Al(s)spotSi
Si(2σ)inst
Si(s)spotPP(2σ)inst
P(s)spot
S
S(2σ)inst
S(s)spot
K
K(2σ)inst
K(s)spot
Ca
Ca(2σ)inst
Ca(s)spot
Ti
Ti(2σ)inst
Ti(s)spotFeFe(2σ)inst
Fe(s)spot
Cu
Cu(2σ)inst
Cu(s)spotMo
Mo(2σ)inst
Mo(s)spotbalance

0.000
1.146
0.000
2.089
0.153
3.105
27.146
0.167
4.108
0.047
0.043
0.081
21.206
0.150
6.641
0.087
0.026
0.077
0.032
0.012
0.014
0.000
0.150
0.000
11.9180.089
1.197
14.803
0.160
3.822
0.000
0.002
0.000
22.672

IP-12

L-1
0.061
0.201
0.106
0.516
0.058
0.802
46.957
0.210
4.859
0.085
0.035
0.023
0.825
0.014
0.642
0.361
0.010
0.573
0.011
0.004
0.001
0.000
0.031
0.000
0.498
0.018
0.247
0.199
0.004
0.166
0.009
0.002
0.016
50.479

0.269
0.261
0.233
7.597
0.165
0.232
39.581
0.227
0.925
0.137
0.024
0.044
2.322
0.031
0.295
4.530
0.057
0.073
0.257
0.022
0.082
0.358
0.102
0.032
2.275
0.034
0.097
0.000
0.003
0.000
0.000
0.002
0.000
42.674

Mo-2
0.000
0.243
0.000
0.552
0.066
0.072
53.778
0.228
2.258
0.000
0.042
0.000
2.969
0.033
3.238
0.296
0.014
0.042
0.030
0.005
0.003
0.126
0.045
0.016
0.856
0.023
0.986
0.061
0.003
0.055
0.000
0.002
0.000
41.333

Mo-4
0.377
0.273
0.327
5.698
0.148
3.146
39.729
0.205
7.880
0.030
0.031
0.052
0.866
0.018
0.515
5.385
0.057
2.497
2.443
0.044
0.892
0.263
0.084
0.171
1.167
0.026
0.576
0.409
0.006
0.235
0.007
0.001
0.004
43.625

Mo-6
0.303
0.307
0.270
6.713
0.167
0.808
36.751
0.204
3.732
0.056
0.024
0.016
3.018
0.037
1.362
5.082
0.058
1.210
2.069
0.043
1.472
0.241
0.091
0.030
2.134
0.031
0.828
0.757
0.009
0.545
0.008
0.001
0.006
42.867

Mo-8

Table 4. Cont.

1.051
0.357
0.493
5.670
0.167
0.968
29.638
0.182
4.311
0.014
0.051
0.024
11.947
0.091
8.311
1.676
0.043
0.058
2.309
0.054
0.387
0.137
0.131
0.120
6.721
0.055
2.205
0.024
0.002
0.008
0.009
0.001
0.006
40.805

OH-1

S-1
1.292
0.317
0.421
7.591
0.187
0.344
32.075
0.216
1.404
0.108
0.024
0.004
3.284
0.041
0.953
3.363
0.056
2.083
3.203
0.062
0.719
0.404
0.135
0.018
4.254
0.045
0.253
0.033
0.002
0.017
0.000
0.002
0.000
44.392

S-2
0.202
0.440
0.350
4.372
0.150
0.426
36.292
0.203
1.954
0.239
0.027
0.161
5.937
0.057
1.410
3.247
0.050
0.498
2.203
0.046
0.473
0.465
0.110
0.325
2.922
0.036
1.337
0.930
0.010
0.640
0.000
0.002
0.000
43.190

S-4
0.595
0.261
0.166
6.134
0.156
0.395
36.938
0.202
2.825
0.071
0.024
0.028
4.012
0.041
3.240
5.932
0.065
1.774
0.888
0.034
0.402
0.168
0.106
0.185
2.792
0.035
1.392
0.032
0.002
0.038
0.034
0.001
0.005
42.405

0.497
0.287
0.038
7.896
0.182
0.254
34.163
0.187
0.809
0.167
0.026
0.044
7.376
0.064
0.765
4.304
0.056
0.105
1.359
0.039
0.696
0.307
0.102
0.036
3.619
0.037
0.222
0.607
0.008
0.097
0.002
0.001
0.002
39.705

SM-3
2.504
0.401
0.410
8.039
0.212
0.424
28.977
0.243
1.918
0.153
0.023
0.067
1.453
0.027
0.552
5.508
0.093
0.334
0.498
0.040
0.085
1.228
0.235
0.027
8.000
0.066
0.907
0.400
0.007
0.218
0.001
0.002
0.002
43.238

SM-4

Y-3
0.646
0.208
0.315
4.013
0.114
0.729
38.218
0.212
0.839
0.057
0.028
0.059
0.323
0.011
0.197
3.039
0.043
0.905
0.729
0.026
0.215
0.232
0.088
0.118
3.543
0.037
1.221
0.116
0.004
0.085
0.000
0.002
0.000
49.085Minerals 2020, 10, 431

252

Minerals 2020, 10, 431

Figure 5. Plots comparing the composition of samples determined by pXRF with those values calculated
from mineral modes: (a) Cu, (b) Mo, (c) Fe, (d) S, (e) Si, (f) Al, (g) K, and (h) Mg. The grey, coarsely
dashed lines show 1:1 equivalency for both methods.

253

Minerals 2020, 10, 431

The degree of correlation between the two methods depends on the particular element. The lightest
and the heaviest elements are systematically overestimated by optical petrography compared to pXRF:
Cu, Mo, Fe, Al, Mg, and Ti (Figure 5a–c,f,h). In contrast, the mid-atomic number elements areunderestimated by optical petrography compared to pXRF: S, Si, K, Ca, and P (Figure 5d,e,g).
The element that shows the closest 1:1 correlation between the two methods is Al (Figure 5f), followedby Si (Figure 5e). The poorest correlations are shown for Cu and Mo (Figure 5a,b). The correlation for
Cu is far worse when the single high-Cu sample (IP-12) is omitted from the ﬁt (Figure 6). The best
trendline ﬁts are for Cu, S, and Al (Figure 5a,d,f). However, the ﬁt is very poor in terms of Cu if sample
IP-12 is omitted (Figure 6). The worst ﬁt is shown by Mo (R2 = −0.06, Figure 5b), followed by Ca(R2 = 0.443), Ti (R2 = 0.4662), and P (R2 = 0.4721). The balances (remaining wt% not attributed to asingle element) are slightly overestimated by mode compared to pXRF by a factor of 1.0623, and with anadequate ﬁt (R2 = 0.7361). Signiﬁcant outliers from these trends are Bu-06: higher Fe and S, but lower
Si by mode than by pXRF, Bi-02: lower K by mode than by pXRF, F-02, higher K by mode than by pXRF,
L-01: higher Cu by mode than by pXRF, and Mo-02: higher Mg by mode than by pXRF.

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4. Discussion
There are multiple factors aﬀecting each method that may bias results. These aﬀect the accuracy
and precision of the mode-based or pXRF-based methods, and the correlation of the two.
4.1. Causes of Poor Correlation
In terms of optical petrography, melanocratic (i.e., maﬁc minerals, sulﬁdes, and oxides) minerals
are typically overestimated compared to leucocratic minerals (i.e., felsic minerals, carbonates,and phosphates; c.f., [27]). Sulﬁde minerals are exclusive hosts for Cu and Mo in the samples.
Similarly, biotite-chlorite and rutile are the major hosts for Mg and Ti, respectively. Biotite, chlorite,

254

Minerals 2020, 10, 431

hematite, and sulﬁdes are also the hosts for Fe. These minerals are all dark-colored or opaque in TPPL
and readily noted. Colorless or light-colored minerals in TPPL such as quartz, muscovite (as sericite),orthoclase, plagioclase, calcite, and apatite collectively host the bulk of Si, K, Ca, and P. The residency of
Al is less clear as it is overestimated by optical petrography (Figure 5f) and is an important constituent(along with K and Si) of biotite-chlorite but is a more signiﬁcant component of the feldpars and sericite.
Human error as systematic visual overestimation of biotite and chlorite during optical petrography
would give a positive bias to the Mg, Fe, and Al calculated by mode; particularly if it was at theexpense of quartz content. The other element with unusual behavior is S in that it is underestimatedby optical petrography. As the only S-bearing phases noted are sulﬁdes, it would be expected that
optical mineralogy would overestimate this element as with Fe, Cu, and Mo. These elements trackwith S (Figure 4), but it is possible another S-bearing phase was overlooked, such as barite, anhydrite,
or gypsum. Barite is not likely as a negligible Ba content was noted by pXRF in all samples. Hypogene
anhydrite occurs in porphyry-type deposits (e.g., [28]), and can be fairly common [5,29]. Another cause
for S underestimation from the mode is misidentifying sulﬁde minerals that are S-rich (chalcopyrite andcovellite) for S-poor phases (bornite, digenite, and chalcocite). However, the most abundant S-bearing
mineral is pyrite, and a similar underestimation is not present with Fe; nor are those two groups ofCu-sulﬁdes easily confused in RPPL. It may be that the pXRF systematically overestimates S content of
the samples and further calibration should have been applied. Systematic overestimation of S (alongwith P, Cu, and Mg) by pXRF has been observed with ﬁeld analysis of soils [30]. However, Roman
Alday et al. [16] found that for samples from the Elatsite porphyry-type deposit, pXRF overestimates
Fe and K, and underestimates Mg compared to more intensive laboratory methods. Matrix eﬀects
on pXRF vary with mineral assemblages and development of calibration standards is reliant on rocktype [23,31]. The poorest correlation of the chalcophile elements is for Mo. Economic grades inporphyry-type deposits range from approximately 0.1 to 0.3 wt% Mo, although lower grades in thepresence of another commodity such as Cu are common [32]. This is a lower grade than Cu due to thehigher valuation of Mo, typically 2× to 3× the per-pound metal price. This grade range approaches
the limit of detection for Mo by most pXRF units in a typical silicate–sulﬁde matrix. The result isthat although molybdenite is fairly distinct and easy to identify in reﬂected light, it may be in trace
concentrations that contribute grades too low to be detected by pXRF. Thus, mode-based estimates ofgrade will likely be higher than those detected by pXRF. The selection of the pXRF analytical spots
may also play a role (see Section 4.4).
4.2. Outliers from Trends
Five samples, in particular, plotted as major outliers from the trends established by the rest of thedata for one or more elements. These are most likely the result of errors in visual estimation of modes
as discussed above. Bu-06 has 2.8 wt% Fe by pXRF and 15 wt% Fe by mode. This is accompanied by asimilar overestimation of S by mode and underestimation of Si by mode. This is most likely due tothe overestimation of pyrite at the expense of quartz. Bi-02 has 8.5 wt% K by pXRF and 4.2 wt% K bymode. This is likely from an underestimation of muscovite (as sericite). F-02 has 1.9 wt% K by pXRF
and 3.5 wt% by mode. The overestimation of potassic phases is likely. Orthoclase is most probableas it has a higher proportion of K to Al than muscovite and there is not a signiﬁcant Al outlier forthis sample. There is no accompanying Fe and/or Mg outlier, so biotite-chlorite is not likely. L-01 has
0.2 wt% Cu by pXRF and 3.2 wt% Cu by mode. This is most likely due to an overestimate of the modal
abundance of chalcopyrite. Mo-02 has 0.3 wt% Mg by pXRF and 1.9 wt% Mg by mode, most likely due
to an overestimation of a maﬁc mineral, such as chlorite. It should be noted, that the element with thehighest instrumental error is Mg (Table 4).
4.3. Assumptions of Mineral Composition
Other than errors in visually estimating the mode of a sample or in misidentifying phases, anothersource of error in calculating the bulk rock composition from the mode is the assumed composition

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of the mineral phases. Chlorite, biotite, and plagioclase all have assumed compositions in termsof solid solution components for the purposes of translating the modes into compositional data.
These assumptions are not supported by any mineral analysis such as electron microprobe analysis.
Minerals such as biotite can vary in composition between diﬀerent deposits and within single deposits.
Magmatic biotite of the host is typically more Fe-rich than the later hydrothermal biotite [6], largely
depending on what alteration zone (i.e., potassic vs. phyllic) they are from (e.g., [33]). There was no
distinction between these parageneses of the same minerals during the determination of the modes.
As chlorite is largely a product of biotite alteration in these deposits, a similar degree of compositionalvariation is expected. Likewise, plagioclase in porphyry deposits varies in composition. For example,
plagioclase from Sierrita ranges from Ab25 to Ab45 [34]. The simpliﬁed estimation of a molar 1:1 ratioof Na:Ca in plagioclase in this study is likely why bulk compositions based on modes overestimates
Ca content. Muscovite and chlorite from the Highland Valley district (that includes the Lornex and
Bethlehem deposits) range in composition 2.2–2.6 Al apfu and 0.5–3.5 Mg apfu, respectively [35].
4.4. Sampling Biases
A further source of disparity between the results of the two methods lies in sampling. Though thethin section is cut from the billet, there may be fairly distinct diﬀerences in mode between a 0.03 mm
thick section and a ~10 mm thick billet. The thin section is essentially a 2D representation of a 3D
rock, whereas the pXRF analysis is an average of three diﬀerent spots on the billet, each 8 mm indiameter (the default aperture size on the Niton XL3t), with an eﬀective sampling depth (d) expressedby Equation (1):d = m/ρ
(1)
where d is the sampling depth from which 99% of the signal is sourced, m is the mass per unit
area sampled and ρ is sample density [23]. For the samples, the least dense abundant mineral isplagioclase (ρalbite = 2.61 g/cm3 ) and the densest is pyrite (ρpyrite = 5.01 g/cm3 ). Denser samples(i.e., sulﬁde-rich) will approach more surﬁcial analyses, whereas sampling depth will be deeper for
samples dominated by felsic minerals. If the billet has compositional variations in the vertical (depth)
sense, then the spot analyses become pseudo-bulk analyses of a cylinder 8 mm in diameter a a depththat is a function of sample density. Similarly, in the horizontal sense, three spots at 50.27 mm2
totals ~150.8 mm2 , whereas the thin section is up to 1196 mm2 in total area. Accounting for 3 mm
margins around the billet, this reduces to 800 mm2 of sample area. Thus, this method in the best case,sampled only ~19% of the sample area and assumes it is representative of the other 21%. Sampling
precision for ﬁner-grained rocks by pXRF has been shown to be ≤5% relative standard deviation,
but much higher for coarser-grained samples [36]. The thin section is a ~60 μm slice oﬀ of the top of
the billet that is subsequently ground down to 30 μm. Sample heterogeneity is typically on a scale
an order of magnitude larger. Diﬀerences in thin section and sample billet surface are negligible incomparison to the other discussed sources of uncertainty. Although most of the rocks in this sample
suite are hypabyssal to volcanic, the presence of phenocrysts and veins/veinlets increases the numberof measurements required to achieve the same precision that would be possible compared to an
un-veined, aphyric sample. The large sample standard deviations across the spot analyses ((s)spot )for many samples illustrates this (Table 4). Samples with a high degree of heterogeneity in terms of
multiple elements are all the Ithaca Peak samples, Be-4, Bu-2, Bu-7, F-3. Mo-6, OH-1, and S-4.
4.5. Applicability of pXRF to Mineral Exploration
Compared to the classic standard of optical petrography, pXRF is faster, more accessible, portable,
and more easily conducted with little training. It is not a replacement for laboratory methods as it cannot
reliably reach the levels of precision and accuracy required to meet standards for mineral resource
reporting in terms of NI 43-101 (National Instrument 43-101 is the Canadian code for the Standardsof Disclosure for Mineral Projects. It sets standards for reporting mineral resource information such

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as grade, tonnage, etc., and what methods are suitable for measuring them to an adequate degree ofcertainty) or JORC (Joint Ore Reserves Committee Code is essentially the Australian equivalent to NI
43-101) guidelines, but with calibration, it can achieve a low but useful level of quantitative certainty(e.g., [16,20,21]). However, for quick decisions regarding unknown samples in the ﬁeld or core shed,semi-quantitative analysis is satisfactory in order to make rough interpretations on whether to collecta sample for more detailed assay/examination. Similarly, pXRF can be used to guide petrographic
examinations in terms of possible (and impossible) phases present as with this study.
Although this study attempts to simulate “ﬁeld” use of pXRF, the analytical conditions for thisstudy were more optimal and controlled than those typical of ﬁeld or even core shed settings. Analyzed
surfaces of the billets were fresh, clean, and ﬂat. Samples from outcrop, ﬂoat, or trenches may have soil,
lichen, or other debris covering the rock. Thick weathering rinds may also be present. Surfaces are
typically irregular. Although conditions for drill core analysis are more controlled, sample geometry is
not ﬂat, but curved. This is especially relevant for smaller diameter drill core such as AQ, BQ, or NQ.
pXRF analysis of reverse circulation (RC) chips suﬀers from especially variable conditions in terms ofsample geometry and open space between the chips. These sources of uncertainty can be mitigatedby cleaning surfaces, preparing ﬂat surfaces, and even pulping and packing them into homogenizedpucks, but all at the cost of signiﬁcant time.
4.6. Directions for Future Research
The certainty of pXRF results would be much improved from non-factory calibration speciﬁc
to porphyry-type samples. On the petrographic side, determination of the variations in mineralcomposition, at least of biotite, chlorite, plagioclase, and muscovite is needed in order to provide
more accuracy in determining a chemical composition from the mode. The use of a scanning electron
microscope with energy-dispersive X-ray spectra would provide suﬃcient accuracy. More detailed
petrographic work, along with cathodoluminescence would help distinguish between diﬀerent
parageneses of minerals and better characterize the range of compositions present (c.f., [29,37]).
Author Contributions: Conceptualization, A.D.V.; methodology, A.D.V.; formal analysis, C.A.G.; investigation,
C.A.G.; resources, A.D.V.; data curation, C.A.G.; writing—original draft preparation, C.A.G.; writing—review andediting, A.D.V. and C.A.G.; supervision, A.D.V.; project administration, A.D.V.; funding acquisition, C.A.G. and
A.D.V. All authors have read and agreed to the published version of the manuscript.
Funding: Funding for Connor Gray was provided by a Bloomsburg University Undergraduate Research, Scholarly,
and Creative Activity (URSCA) award.
Acknowledgments: Samples were collected by Peter H. Kirwin and John Ray. Funding for the Leica microscope
was provided by the Bloomsburg University College of Science and Technology Dean’s oﬃce via Troy Prutzman
and Robert Aronstam. I, Connor Gray, want to personally thank my undergraduate research advisor Adrian Van
Rythoven for the opportunity to work on this project. Two anonymous reviewers and the editorial staﬀ at Minerals
provided comments that improved the quality of the manuscript.
Conﬂicts of Interest: The authors declare no conﬂict of interest.

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